Keay Thesis 1998

The Geological Evolution of the Cyclades, Greece:

Constraints from

SHRIMP U-Pb Geochronology.

by

Sue Keay

A thesis submitted for the degree of Doctor of Philosophy of the Australian National University.

1998

iii

Declaration

This thesis contains no material which has been accepted for the award of any other degree or diploma in any university or other institution, and to the best of my

knowledge contains no material previously published or written by another person, except where due reference is made in the text.

Sue Keay 1998

iv

"few dates can be given with any certainty. I regret this, for it seems to me that in history, which Napolean defined as an agreed fable, dates have a reassuring firmness about them, and for my part, I like to have plenty around. I am grateful to those who salvage dates for us, and even to those who make a gallant attempt at it, such as Dr John Lightfoot who, in 1642, affirmed that the world "was created by the Trinity on

October 23rd 4004BC at 9 o'clock in the morning." The Doctor may have been wrong, but he had, I feel, the right approach to history!".

Eric Forbes-Boyd

Aegean Quest A Search for Venetian Greece

1970

v

Acknowledgements

vi

Abstract

vii

Table of Contents

Acknowledgments These acknowledgments are unashamedly long. One-seventh of my life has been

devoted to this thesis and I have a lot of people to thank for sharing this time with me. First I’d like to thank my supervisors, Bill Compston and Gordon Lister, for their

constant support over the last few years and, in particular, the last trying months. Bill speedily read material for me at incredibly short notice and without complaint while Gordon’s unquenchable thirst for knowledge and boundless enthusiasm for geology never cease to amaze me. During the course of my PhD I have enjoyed the friendship and hospitality of Elizabeth Compston and Pam Lister, thank you.

I would never have submitted a PhD thesis if it wasn’t for the stalwart support of David Ellis. Dave intervened on my behalf at crucial times during my PhD using his amazing powers of lateral thinking to ensure I got a fair go.

My Honours supervisor, Bill Collins, continues to offer me encouragement, despite the distance between us, and I have been strongly influenced by Bill’s ever-questioning approach to science and his fearlessness in challenging established dogma.

My PhD adviser Ian Buick has done his best to help me appreciate metamorphic petrology (almost successfully) while also providing tremendous support, much needed motivation and a 24-hour-return reading service!

I don’t know where I would have been without my de-facto PhD adviser, Ian Fitzsimons over the last two months (somewhere on another planet perhaps). Ian generously gave up his time to offer me critical comments and invaluable advice when I needed it most and it seems unlikely I’ll ever be able to return the favour.

I would like to thank most of the people at RSES (and beyond!) for their friendship/help/encouragement at one time or another - I think I owe everyone a beer so looks like the keg’s on at my place! I was humbled by the number of people who generously gave up their time in order to help me complete this work, especially those who offered their proof-reading services in the last mad days before submission. My thanks go to Eleanor Dixon, Mark Jellinek, Alfredo Camacho, John Mavrogenes, Dan Zwartz, Graham Hughes, Richard Armstrong, Rob Scott, Lance Black, Keith Sircombe, Ken Lawrie, Chuck Magee, Stefan Klemme, Ulrike Troitzsch for titanite references ASAP, several others whose offers I couldn’t take up, I even got comments from Utah thanks to Dave Dinter. My appendices (and sanity) are greatly indebted to the help and moral support of Kevin Fleming. I wish I could have used one of your quotes in my thesis Kevin, but I wasn’t sure other people would understand the supposed likenesses between my thesis and the struggle of the good forces of the universe against evil - nor the supposed similarity between myself and Captain Janeway. Megan Hough and Sarah Vaughan from Monash University helped me enormously with drafting figures, and I appreciate Gordon allowing them the time to help.

The RSES is blessed with an outstanding array of technical and administrative staff and I have called on most of them for assistance in times of difficulty (technical and otherwise!). Anne Gillard dealt with the traumas of my arrest in Greece with her usual remarkable cool efficiency, and she and Clementine Krayshek have dealt with many phone calls for advice on thesis production in the last month.

While the RSES is well-equipped in terms of personnel, equipment and a general ambience conducive to scientific research, I found myself suffering from an almost terminal lack of enthusiasm 2/3rds of the way through my project. Two things helped me to pull through this difficult time: a visit to Mainz university in Germany and a

Penrose conference in Crete where Mark Brandon interrogated me mercilessly about my work and why I was doing it and Steve Reddy asked endless questions about zircons.

While travelling through Greece and Germany I have enjoyed the generous hospitality of many people, in particular I would like to thank Nicos Katsaris and Nicole Eder of Naxos, for food/lodging/transport and the best citron in the Cyclades. Rolf Claesson pointed out the historical significance of Naxos (as well as the phone number of the best lawyer during my arrest), while Samantha Barr and John Tarney showed me around the sights of the Rhodope complex. In Germany, Professor Stefan Dürr took me in as part of his own family, while Joerk Jarick and Petra introduced me to apple wine.

My time in Canberra has been greatly enriched by my association with the Commonwealth Territories Division of the Geological Society of Australia. Through the society I have met a broad spectrum of geologists and I will fondly remember evenings spent listening to Mike Rickard and Larry Harrington explaining the development of ideas in Australian geology and dobbing in all the academics who were slow to accept the theory of Plate Tectonics.

I have spent many pleasant hours watching the sunset over Lake Burley Griffen from the grounds of Old Canberra House, thanks largely to the company of Richard Wysoczanski, Rich Armstrong, Mark Fanning, Jim Dunlap, Paul Hoskin, Paul Johnston, Melita Keywood, Dylan, Dave John Brown, Claudine Stirling and Geoff Fraser. Keith Sircombe, Michael Wingate and Paul Hoskin have shared the peculiariarities of being a SHRIMP student in the RSES dungeon. Inside and outside drinking hours Mark and I have enjoyed the friendship of Corine Davids and Tony Doulgeris.

The beauty of having two supervisors is that I have also had two universities to choose from - ANU and Monash. At Monash University I have enjoyed the pleasurable company of; Jodie Miller, Caroline Read, Marnie Forster, Terence Barr, Tim Rawlings, Tyler MacCready, John Miller, Ed Curl, Caroline Venn and Mandy Raouzaios.

The evolution of my thesis (and life) has not always been ideal and I have cried on many of the following people’s shoulders (for which I would like to offer my gratitude and a spare box of tissues): my good friend Alfredo Camacho is always there for me no matter what trivial thing I’m upset about; Steve “regolith” Hill has also picked me up after numerous falls; Robin Maier, the RSES students’ “Mom”; my running partner and SHRIMP-fixer extraordinaire John “Fozzie” Foster; John “the prince” Mya; Shane “landlord” Paxton; Mark “are you done yet?” Jellinek (it’s Mai Tai time!) and Kevin the hippie tea king. My life has been immeasurably brightened by the friendships of three people who never seem to go anywhere without a smile, Eleanor Dixon, Erica Hendy and my writing-up partner Monica Handler.

My family has been incredibly forbearing during the course of my thesis. My parents spent their fortieth wedding anniversary helping me put things together, my brother Lindsay ran around providing technical support while my mother-in-law sent food from Newcastle to ensure I got fed. I appreciate all the sacrifices my family have made and the completion of this thesis is largely due to their help. None of this extraordainary odyssey would have been possible without the help of my husband Mark, who, despite his accountant’s financial advice to, “get your wife a real job”, continues to encourage and support me in any endeavour and who tolerates all my eccentricities.

ABSTRACT

The Alpine-Himalayan mountain chain formed by the collision of continents and is

one of the most dramatic manifestations of plate interaction on the Earth’s surface. The

Cyclades are a group of islands located between the converging African and Eurasian

plates and their evolution is inextricably linked with that of the Alpine-Himalayan

orogenic belt (from ~140 Ma to present). The Cyclades consist of three main lithological

groups: two of these, the Mesozoic Series and pre-Mesozoic Basement, have undergone

high-P metamorphism associated with the collision of Africa and Eurasia, the third, the

Upper Unit, has not experienced the effects of the Alpine orogeny. A SHRIMP

(Sensitive High Resolution Ion MicroProbe) geochronological investigation of U,Th-

bearing minerals with high closure temperature from both magmatic and metamorphic

rocks has shed light on both the Alpine polymetamorphic history of the Cyclades and

also, through the use of zircon provenance ages, the earlier tectonic history of the region

prior to the collision of Africa and Eurasia.

In the Cyclades, metamorphism associated with the Alpine orogeny has resulted in

the development of new hydrothermal zircon growth, with zircon ages reflecting the

timing of multiple episodes of fluid infiltration during sub-solidus metamorphism. Fluid

flow must have occurred in response to some trigger, most likely tectonic activity, and so

zircon ages, in conjunction with other data, can be used in the construction of

metamorphic pressure-temperature-time paths for the area. New zircon growth does not

occur in all lithologies but, where it does occur, there is a consistency in zircon ages from

different rock units and even from different islands. This reproducible and previously

unreported complexity in zircon ages suggests that widespread formation of zircon occurs

under a range of metamorphic conditions associated with fluid infiltration. The absence

of new zircon growth in some lithologies can be related to their permeability and

susceptibility to interaction with fluids.

Populations of zircon ages in Cycladic samples from detrital and inherited grains

appear to correlate well with periods of tectonic activity related to movements of the

African and Eurasian plates during and prior to the Alpine orogeny. As the Cyclades

formed part of the northern margin of Gondwana, this new compilation of zircon ages for

the Cyclades provides a reference database that characterises the source of this crustal

material. SHRIMP U-Pb dating of granitic orthogneisses has shown that the Basement

units of the Cyclades record extensive granitoid formation during the late stages of the

Variscan orogeny (330-300 Ma) associated with the collision of Gondwana and Laurasia.

Subsequent rifting along the northern margin of Gondwana during the Triassic-Jurassic is

recorded by magmatic zircons in the sediments of the Cycladic Series rocks, marking a

time of active magmatic activity and associated sedimentation in the region that is common

to many areas of the Alpine chain. Separation of microcontinental blocks from the

northern margin of Africa and their movement towards the southern margin of Eurasia

resulted in the formation of rifts floored by oceanic crust. Late Cretaceous (~ 75 Ma)

zircons from a high-P metaophiolite of the Cyclades suggests that oceanic basins were in

existence at this time during the convergence of African and Eurasia. The timing of

events associated with the Alpine orogeny is preserved by a range of Cretaceous and

Tertiary-aged metamorphic zircon overgrowths. The youngest of these ages on the island

of Naxos reflects the timing of partial melting and associated Barrovian metamorphism at

ca . 18 Ma, followed by late stage fluid movement associated with extensional shearing at

ca . 13 Ma. The voluminous intrusion of small granitoid bodies on Naxos closely follows

the timing of peak metamorphism, occurring mainly at ca . 12 Ma.

Extracting age information from the complicated zircons of the Cyclades has only

been possible using the microanalytical capability of the Australian National University’s

SHRIMP ion microprobe. SHRIMP provides high precision measurement of isotopic

species under conditions of both high mass resolution (to exclude molecular interferences)

and high spatial resolution, allowing analysis of areas within mineral grains which are

less than 20 microns in diameter and a few microns deep. The application of this

technique to a relatively young terrane has enabled the recognition of multiple episodes of

new zircon growth at metamorphic grades where zircon is generally considered to remain

inert.

vii

TABLE OF CONTENTS

1. INTRODUCTION. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1

1.1 AIMS OF THIS STUDY ..........................................................................................................3

1.2 TECTONIC SETTING OF THE CYCLADES ...................................................................................3

1.3 REGIONAL GEOLOGY .........................................................................................................7

1.4 GEOLOGY OF THE CYCLADES ...............................................................................................8

1.5 METAMORPHIC HISTORY OF THE CYCLADES.......................................................................... 10

1.6 STRUCTURAL DEVELOPMENT IN THE CYCLADES .................................................................... 11

1.7 LOCAL GEOLOGY............................................................................................................. 14

1.7.1 Naxos...................................................................................................................... 14

1.7.2 Paros...................................................................................................................... 16

1.7.3 Ios .......................................................................................................................... 17

1.7.4 Syros ...................................................................................................................... 18

1.7.5 Sifnos...................................................................................................................... 19

1.7.6 Sikinos.................................................................................................................... 21

1.7.7 Folegandros............................................................................................................. 22

1.7.8 Tinos ...................................................................................................................... 23

1.8 SHRIMP U, TH-PB GEOCHRONOLOGY OF ACCESSORY MIMERALS ............................................ 24

1.8.1 Closure Temperature (TC) .......................................................................................... 25

1.8.2 Application of Geochronology to Metamorphic Processes............................................... 26

1.9 ZIRCON.......................................................................................................................... 28

1.9.1 Metamorphic Zircon................................................................................................. 30

1.9.2 Th/U Chemistry of Zircons......................................................................................... 33

1.10 MONAZITE.................................................................................................................... 35

1.10.1 Metamorphic Monazite ........................................................................................... 36

1.11 TITANITE...................................................................................................................... 38

1.11.1 Metamorphic Titanite.............................................................................................. 39

1.12 STABLE ISOTOPES........................................................................................................... 40

1.13 APPLICATION OF SHRIMP DATING TO THE CYCLADES.......................................................... 42

1.13.1 Zircon Inheritance Patterns...................................................................................... 43

1.14 SAMPLE SELECTION........................................................................................................ 46

2. PRE-CARBONIFEROUS EVOLUTION OF THE CYCLADES. . . . . . . . . . . . . . . . . . . . . . . . 4 9

2.1 INTRODUCTION................................................................................................................ 49

2.2 GEOLOGIC SETTING ......................................................................................................... 51

2.3 PREVIOUS GEOCHRONOLOGY............................................................................................. 52

2.4 U-PB ANALYTICAL RESULTS.............................................................................................. 53

2.5 DISCUSSION .................................................................................................................... 56

2.6 SYNTHESIS ..................................................................................................................... 60

viii

3. PERMO-CARBONIFEROUS GEOLOGICAL EVOLUTION OF THE CYCLADES 6 1

3.1 INTRODUCTION................................................................................................................ 61

3.2 GEOLOGICAL BACKGROUND .............................................................................................. 62


3.4 SHRIMP U-PB RESULTS................................................................................................... 64

3.4.1 Ios .......................................................................................................................... 64

3.4.2 Paros...................................................................................................................... 72

3.4.3 Sikinos.................................................................................................................... 75

3.4.4 Naxos...................................................................................................................... 76

3.5 COMBINED PERMO-CARBONIFEROUS RESULTS....................................................................... 84

3.6 DISCUSSION .................................................................................................................... 85

3.6.1 Confirmation of the Existence of Pre-Mesozoic Basement................................................ 85

3.6.2 The Timing of Pre-Alpine Metamorphism M0................................................................. 86

3.6.3 Complications within the Naxos core............................................................................ 86

3.6.4 Correlations with North Africa ................................................................................... 87

3.6.5 Correlations with the Menderes Massif, Turkey............................................................ 88

3.6.6 Correlations with the Pelagonian Zone, Internal Hellenides, Greece ................................ 88

3.6.7 Correlations with the External Hellenides, Crete........................................................... 89

3.6.8 Tectonic Implications of Age Data ............................................................................... 89

3.7 SYNTHESIS ..................................................................................................................... 90

4. TRIASSIC/JURASSIC GEOLOGICAL EVOLUTION OF THE CYCLADES . . . . . . . 9 1

4.1 INTRODUCTION................................................................................................................ 91

4.2 GEOLOGICAL BACKGROUND .............................................................................................. 92


4.4 SHRIMP U-PB RESULTS FOR THE VARI GNEISS ..................................................................... 95

4.5 DEPOSITIONAL AGE AND PROVENANCE OF SEDIMENTS IN THE CYCLADES .................................. 96

4.5.1 Syros ...................................................................................................................... 97

4.5.2 Naxos...................................................................................................................... 99

4.5.3 Ios ........................................................................................................................ 111

4.5.4 Folegandros........................................................................................................... 114

4.5.5 Sikinos.................................................................................................................. 115

4.5.6 Sifnos.................................................................................................................... 116

4.5.7 Combined Triassic-Jurassic U-Pb Zircon Ages............................................................. 120

4.6 DISCUSSION .................................................................................................................. 121

4.6.1 Age of the Vari Gneiss, Syros.................................................................................... 121

4.6.2 Age of Series Rocks................................................................................................. 121

4.6.3 Correlations with the Menderes Massif, Turkey.......................................................... 123

4.6.4 Correlations with the Pelagonian Zone, Internal Hellenides, Greece .............................. 125

4.6.5 Correlations with the External Hellenides................................................................... 125

4.6.6 Tectonic Implications of SHRIMP Ages ...................................................................... 126

4.7 SYNTHESIS ................................................................................................................... 129

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5. CRETACEOUS GEOLOGICAL EVOLUTION OF THE CYCLADES . . . . . . . . . . . . . . 1 2 9

5.1 INTRODUCTION.............................................................................................................. 129

5.2 GEOLOGICAL BACKGROUND ............................................................................................ 130

5.3 PREVIOUS GEOCHRONOLOGY........................................................................................... 131

5.4 DATING OF THE SYROS OPHIOLITE .................................................................................... 133

5.4.1 Zircon Geochemistry .............................................................................................. 135

5.5 CRETACEOUS-AGED ZIRCON OVERGROWTHS....................................................................... 139

5.6 DISCUSSION .................................................................................................................. 142

5.6.1 Dating Ophiolite Formation ...................................................................................... 142

5.6.2 Similarities Between Upper Unit and Series Rocks ....................................................... 143

5.6.3 Correlations Between Cycladic High-P Metaophiolite Units.......................................... 143

5.6.4 Constraints on the Timing of High-P Metamorphism (M1)............................................... 144

5.6.5 Correlations with the Menderes Massif ..................................................................... 144

5.6.6 Correlations with the Pelagonian Zone, Internal Hellenides, Greece .............................. 145

5.6.7 Correlations with the External Hellenides................................................................... 145

5.6.8 Tectonic Implications.............................................................................................. 146

5.7 SYNTHESIS ................................................................................................................... 149

6. TERTIARY METAMORPHIC EVOLUTION OF THE CYCLADES. . . . . . . . . . . . . . . . 1 5 1

6.1 INTRODUCTION.............................................................................................................. 151


6.2.1 Naxos.................................................................................................................... 152

6.2.2 Sifnos.................................................................................................................... 154

6.2.3 Ios ........................................................................................................................ 155

6.2.4 Other Cycladic Islands ............................................................................................ 156

6.3 EVIDENCE OF FLUID INFILTRATION IN CYCLADIC ROCKS........................................ 156

6.3.1 Naxos.................................................................................................................... 156

6.3.2 Sifnos.................................................................................................................... 157

6.3.3 Ios ........................................................................................................................ 158

6.3.4 Other Cycladic Islands ............................................................................................ 158

6.4 COMPILATION OF EVENTS AND AGE DATA FOR THE CYCLADES............................................... 158

6.4.1 Sample Selection .................................................................................................... 162

6.5 SHRIMP U-TH-PB ZIRCON RESULTS................................................................................. 164

6.5.1 Naxos.................................................................................................................... 164

6.5.2 Sifnos.................................................................................................................... 170

6.5.3 Ios ........................................................................................................................ 171

6.5.4 Combined Metamorphic Zircon Age Results For Cyclades ............................................ 172

6.6 SHRIMP TH-PB DATING OF MONAZITE............................................................................. 175

6.6.1 Comparison of Monazite Ages................................................................................... 178

6.7 SHRIMP U-PB DATING OF TITANITE ................................................................................ 179

6.7.1 Comparison of Titanite Ages..................................................................................... 182

6.8 CORRECTIONS FOR ISOTOPE DISEQUILIBRIUM....................................................................... 182

6.9 STABLE ISOTOPE RESULTS ............................................................................................... 184

6.10 METAMORPHIC FLUID COMPOSITION ............................................................................... 186

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6.11 THE ROLE OF FLUIDS.................................................................................................... 188

6.12 GEOLOGICAL SIGNIFICANCE OF METAMORPHIC ZIRCON....................................................... 189

6.13 U-PB AGES OF METAMORPHIC MINERALS.......................................................................... 190

6.14 RELATING AGES TO P-T-T PATHS ................................................................................... 193

6.15 TECTONIC IMPLICATIONS ............................................................................................... 197

6.16 DIRECTIONS FOR FUTURE RESEARCH............................................................................... 197

6.17 SYNTHESIS................................................................................................................. 198

7. MIOCENE MAGMATIC EVOLUTION OF THE CYCLADES. . . . . . . . . . . . . . . . . . . . . . . 1 9 9

7.1 INTRODUCTION.............................................................................................................. 199


7.2.1 Naxos.................................................................................................................... 200

7.2.2 Tinos .................................................................................................................... 201

7.3 SHRIMP U-TH-PB RESULTS............................................................................................ 202

7.3.1 Zircons ................................................................................................................. 202

7.3.2 Monazite............................................................................................................... 208

7.3.3 Titanite ................................................................................................................. 210

7.3.4 Corrections for Isotope Disequilibrium....................................................................... 212

7.3.5 Combined Zircon/Monazite/Titanite Intrusion Ages..................................................... 214

7.4 DISCUSSION .................................................................................................................. 214

7.4.1 The Effects of Post-Igneous Pb loss............................................................................. 214

7.4.2 Crystallisation/Emplacement Ages............................................................................ 215

7.4.3 Relationship Between Metamorphism and Magmatism.................................................. 217

7.4.4 Comparison to Surrounding Areas.............................................................................. 218

7.4.5 Tectonic Implications.............................................................................................. 218

7.5 SYNTHESIS ................................................................................................................... 220

8. SYNTHESIS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2 2 1

8.1 MESOPROTEROZOIC-ARCHAEAN........................................................................................ 222

8.2 NEOPROTEROZOIC.......................................................................................................... 222

8.3 EARLY PALAEOZOIC....................................................................................................... 223

8.4 PERMO-CARBONIFEROUS................................................................................................. 223

8.5 TRIASSIC-JURASSIC......................................................................................................... 224

8.6 CRETACEOUS ................................................................................................................ 224

8.7 TERTIARY METAMORPHIC EVOLUTION............................................................................... 225

8.8 MIOCENE MAGMATIC ACTIVITY....................................................................................... 225

APPENDICES

A : PUBLISHED/SUBMITTED WORKS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2 2 7

A 1. SUBMITTED PAPERS...................................................................................................... 227

A 2. OTHER PUBLICATIONS................................................................................................... 227

A 3. CONFERENCE ABSTRACTS............................................................................................... 228

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B : SAMPLE LOCATION. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2 3 1

C : SAMPLE PREPARATION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2 3 3

C 1. ROCK CRUSHING.......................................................................................................... 233

C 2. MINERAL SEPARATION.................................................................................................. 233

C 3. SHRIMP MOUNT PREPARATION..................................................................................... 233

C 4. SHRIMP MOUNT IMAGING ........................................................................................... 234

D : ANALYTICAL PROCEDURE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2 3 5

D 1. RADIOACTIVE DECAY................................................................................................... 235

D 2. U-TH-PB GEOCHRONOLOGY........................................................................................... 236

D 3. SECONDARY ION MASS SPECTROMETRY............................................................................ 238

D 4. SENSITIVE HIGH MASS RESOLUTION ION MICROPROBE (SHRIMP) ........................................ 239

D 5. SHRIMP DATA COLEECTION......................................................................................... 241

D 6. SHRIMP DATA REDUCTION.......................................................................................... 244

D 6.1 SHRIMP Standards................................................................................................. 244

D 6.2 Hydride Interferences............................................................................................. 245

D 6.3 Calculation of inter-element ratios............................................................................ 246

D 6.4 Common Pb corrections........................................................................................... 251

D 6.5 Common Pb composition.......................................................................................... 252

D 6.6 Calculation of Radiogenic Isotope Ratios.................................................................... 256

D 7. ISOTOPIC DISEQUILIBRIUM ............................................................................................. 258

D 8. SHRIMP ERROR ANALYSIS............................................................................................ 262

D 8.1 Error on Titanite Ages............................................................................................. 264

D 9. MIXTURE MODELLING.................................................................................................. 264

D 10. STATISTICAL TESTS ON AGE DATA................................................................................. 264

D 10.1 Test of Adequacy.................................................................................................. 264

D 10.2 Significant Differences in Ages................................................................................ 265

D 11. KALEIDAGRAPHTM PROGRAMS....................................................................................... 265

E : U-TH-PB ANALYTICAL RESULTS. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2 7 1

E 1. ZIRCON U-PB ANALYTICAL RESULTS................................................................................ 273

E 1.1 IO9403 Ios Orthogneiss (Z1978, 97759)....................................................................... 273

E 1.2 IO9404 Ios Orthogneiss (Z1978, 97760)....................................................................... 274

E 1.3 89640 Ios Orthogneiss (Z2405, 89640) ........................................................................ 274

E 1.4 IO9607 Ios Leucogneiss (Z2665, 97761)...................................................................... 275

E 1.5 IO9606 Ios Garnet Mica Schist (Z2665, 97762)............................................................ 275

E 1.6 IO9609 Ios Garnet Mica Schist (Z2665, 97763)............................................................ 276

E 1.7 PA9606 Paros Orthogneiss (Z2644, 97764).................................................................. 276

E 1.8 PA9601 Paros Orthogneiss (Z2665, 97765).................................................................. 277

E 1.9 SK9601 Sikinos Orthogneiss (Z2633, 97766)................................................................ 277

E 1.10 NX9314 Naxos Layered Acid Gneiss (Z1889, 97767)................................................... 278

E 1.11 NX9485 Naxos Layered Acid Gneiss (Z2645, 97768)................................................... 278

E 1.12 NX9315 Naxos Leucogneiss (Z2264, 97769) ............................................................... 279

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E 1.15 NX94103 Naxos Migmatite (Z2153, 97772)................................................................ 283

E 1.16 NX9638 Naxos Migmatite (Z2665, 97773).................................................................. 284

E 1.17 NX9637 Melt Pod Naxos Migmatite (Z2782, 97774).................................................... 285

E 1.18 NX9451 Naxos Quartzite (Z2156, 97775)................................................................... 286

E 1.19 NX9481 Naxos Quartzite (Z2217, 97776)................................................................... 287

E 1.20 SY9603 Syros “Vari” Orthogneiss (Z2665, 97777) ...................................................... 288

E 1.21 89646 Syros Quartzite (Z2405, 89646) ...................................................................... 289

E 1.22 SY9630 Syros Schist (Z2644, 97778)......................................................................... 290

E 1.23 NX9461 Naxos Calc-silicate (Z2298, 97779).............................................................. 290


E 1.25 NX94112 Naxos Calc-silicate (Z2298, 97800) ............................................................ 291

E 1.26 NX9464 Naxos Calc-Silicate (Z2038, 97782).............................................................. 292



E 1.29 NX9490 Naxos Pelite (Z2264, 97781)........................................................................ 297

E 1.30 NX94106 Naxos Pelite (Z2298, 97785)...................................................................... 299

E 1.31 Ios Glaucophane Schist (Z2405, 89639) .................................................................... 300

E 1.32 IO9615 Garnet-Glaucophane Schist (Z2644, 97786) ................................................... 301

E 1.33 90346 Ios Quartz-Phengite Schist (Z2405, 90346)........................................................ 302

E 1.34 FL9602 Folegandros Pelite (Z2633, 97787) ............................................................... 303

E 1.35 SK9603 Sikinos Metabasic Schist (Z2633, 97788)...................................................... 304

E 1.36 SIF9345 Sifnos Calc-silicate (Z2363, 97789).............................................................. 305

E 1.37 89642 Syros Retrogressed Eclogite (Z2405, 89642) ..................................................... 306

E 1.38 NX9301 Naxos I-type Granodiorite (Z1870, 97790)..................................................... 307

E 1.39 NX9303 Naxos Fractionated I-type Granite (Z2298, 97791).......................................... 308

E 1.40 NX9470 Naxos I-type Granitoid (Z2613, 97792).......................................................... 308

E 1.41 NX9446 Naxos S-type Granite (Z2613, Z2644, 97793) ................................................. 310

E 1.42 TIN9603 Tinos S-type Granite (Z2665, 97794) ........................................................... 310

E 2. MONAZITE U-TH-PB ANALYTICAL RESULTS............................................................. 312

E 2.1 NX9637 Naxos Melt Pod Naxos Migmatite (Z2922, 97774)............................................. 312

E 2.2 NX94103 Naxos Migmatite (Z2922, 97772).................................................................. 312

E 2.3 NX9315 Naxos Leucogneiss (Z2922, 97769)................................................................. 313

E 2.4 NX9320 Naxos Leucogneiss (Z2922, 97771)................................................................. 313

E 2.5 NX9438 Naxos Pegmatite (Z2301, 97798).................................................................... 314

E 2.6 NX9439 Naxos S-type Granite (Z2037, 97795) ............................................................. 314



E 3. TITANITE U-PB ANALYTICAL RESULTS .................................................................... 317



E 3.3 NX9435 Naxos Amphibolite (Z2265, 97799)................................................................. 318

E 3.4 NX9301 Naxos I-type Granodiorite (Z1858, Z2313, 97790) ............................................ 319

E 3.5 NX9303 Naxos Fractionated I-Type Granite (Z2313, 97791) .......................................... 320

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3 2 1

Chapter 1 1

1. INTRODUCTION

The convergence of Africa and Eurasia caused intense deformation of Greece and

Turkey in the core of the Alpine orogen, adding complexity to reconstructions of the pre-

Alpine and Alpine evolution of the Aegean region. Scattered outcrops of this suture zone

between the two main continental plates, such as the Cyclades (Figure 1-1), have been

exposed by exhumation along low-angle detachments during late and post-Alpine

extension. The Cyclades preserve spectacular outcrops of blueschist and eclogite-facies

rocks representing a mixed continental and oceanic suture zone produced by a complex

history of interactions between the African and Eurasian plates throughout the

Phanerozoic.

GREECE

TURKEY

C Y C L A D E S

AEGEAN SEA

Figure 1-1: The Cyclades are a group of islands located at the junction between Europe and Africa, mid-way between Greece and Turkey. The extent of the Attic-Cycladic Massif is delineated by a dashed line onthis topographic/bathymetric map of the Cyclades, which was generated using a program developed at theResearch School of Earth Sciences by Jean Braun.

The geological evolution of the Cyclades is hard to constrain due to a number of

outstanding problems in the interpretation of the Alpine history of the Aegean.

These include :

1) the number of oceans that disappeared during convergence of Africa and Eurasia;

2) the time at which these oceans were formed and then destroyed by subduction;

2 Introduction3) the number and location of subduction zones operating;

4) the age and provenance of continental crust caught up in the collisional process;

5) the metamorphic evolution of the Aegean region;

6) the timing and cause of extension and exhumation of high-P rocks; and

7) the relationship of magma generation to tectonic processes.

Radiometric age determinations reflect the time at which a rock or mineral became

closed to chemical exchange of the isotopic species being measured. Isotopes in different

minerals are known to respond in characteristic ways to the influence of temperature,

reflecting their diffusivities in different crystal lattices. The temperature at which an

isotope becomes effectively a closed system to isotopic exchange is known as its closure

temperature (Dodson, 1973). With the advent of new high-precision dating techniques,

capable of dating a range of minerals with different closure temperatures, time constraints

can now be placed on the operation of tectonic processes. This information forms an

integral part in the construction of pressure-temperature-time (P-T-t) paths that are used to

summarise the tectonic evolution of crustal segments. The success of this approach is

limited by the extent to which the time of mineral closure can be correlated with particular

geological processes or events. For this reason, the derivation of geologically meaningful

ages through radiometric dating relies on integrating these ages with evidence from other

disciplines before applying them to constrain the timing of tectonic processes.

In order to date the operation of tectonic processes, it is crucial to select an isotopic

system where the uncertainty of age determination is considerably less than the duration

of the process (Zeitler, 1989). This is most easily achieved in young metamorphic

terranes where age uncertainties are relatively small. The Cyclades are an ideal setting for

the investigation of tectonic processes, not only because of their importance in

understanding the history of the Alpine orogen, but also because of their relatively young

age. To generate a geochronological framework for the evolution of the Cyclades, this

study combines a range of new geochronological data with existing knowledge about the

structural, metamorphic and igneous geology of the area. The purpose of this endeavour

is two-fold: to reconstruct the tectonic history of the Cyclades, and to test the application

of SHRIMP U-Pb geochronology to a complicated young terrane.

Previous geochronological and petrological studies of the Cyclades have focussed

on the low temperature part of the P-T-t history. No mineral phase can record the

complete thermal history of a terrane and so to construct a geological time-scale for an

area, a combination of different dating techniques is required. U, Th-Pb geochronology

can be applied to minerals that have relatively high closure temperatures compared to most

other dating techniques, such as K-Ar or Rb-Sr. This study has concentrated on the

application of U, Th-Pb geochronology to the Cyclades because it constrains the high

Chapter 1 3temperature history of the area and because there are no U, Th-Pb datasets available for

the metamorphic history of the Cyclades. This work complements the large body of

existing data concerning the lower temperature history (Altherr et al., 1979; Andriessen et

al., 1979; Wijbrans and McDougall, 1988; Wijbrans et al., 1990; Wijbrans et al., 1993;

Baldwin, 1996; Baldwin and Lister, 1998). In addition, the U, Th-Pb system is likely to

preserve evidence of early events, partly obscured by the Alpine orogeny. To facilitate U,

Th-Pb dating, the Australian National University’s SHRIMP ion microprobe was used to

analyse individual U-bearing minerals. SHRIMP allows within-grain analysis and its

high spatial resolution can reveal internal structures in the isotopic composition of

minerals related to their complicated growth histories.

1.1 Aims of this Study

Multiple metamorphic episodes have been recognised in the Cyclades but the timing

and duration of these episodes is not known in detail. To separate the operation of closely

spaced metamorphic events such as those which might be expected associated with

thermal pulses into the crust, metamorphic minerals from a young terrane, such as the

Cyclades, must be dated at high precision. The unique analytical capabilities of the

SHRIMP ion microprobe allows the early history of crustal precursors to be investigated,

as well as allowing the dating of mineral growth associated with metamorphism. Using

this capability, the present study aimed to constrain not only the Alpine evolution of the

Cyclades but also the pre-Alpine history.

1.2 Tectonic Setting of the Cyclades

The Alpine-Himalayan mountain chain is one of the most dramatic manifestations of

plate interaction on the Earth’s surface (Figure 1-2). A relatively young orogenic belt, it

has a length along strike of over 15000 km. It was formed by a series of collisions

between continental blocks of different character. The Cyclades were located somewhere

between the two main colliding plates, Africa and Eurasia, throughout the Alpine

orogeny. These islands form part of the Apulian-Anatolian or Turkish plate, which

consists of complexly dissected crustal material accreted onto the southern edge of the

Eurasian plate (Smith and Woodcock, 1982). Mediterranean sea-floor, which flanks the

northern margin of the African plate, has been subducted beneath the Turkish plate along

the Hellenic trench (presently located south of Crete) since late Miocene time

(Jacobshagen et al., 1978; Burchfiel, 1980; Robertson and Dixon, 1984). This area

underwent approximately 200-400 km shortening during southward migration of the

Hellenic subduction zone (Schermer et al., 1990). However, despite its overall

convergent setting, the Aegean crust has actually been stretched by a factor of two since

the Miocene (McKenzie, 1977) as a result of transverse plate movements between the

colliding Eurasian, African, Anatolian and Arabian plates (Kempler, 1994) (Figure 1-3).

4 IntroductionE

arly

Pal

aeoz

oic

Oro

geni

c B

elts

:L

ate

Pala

eozo

icM

esoz

oic

Cai

nozo

ic

AL

PIN

E-H

IMA

LA

YA

Figure 1-2: World-map showing the distribution of various Phanerozoic orogenic belts (adapted fromSmith et al., 1981) Unshaded continental areas represent Precambrian shields or cover sediments. TheCyclades form part of the Cainozoic Alpine orogenic belt which extends from western Europe to Asia.Fragments of the Late Palaeozoic Variscan mountain chain occur in western Europe, North Africa andNorth America.

Chapter 1 5Whilst Africa and Europe are still converging at a rate of ~ 2 cm per year (Le Pichon

et al., 1988) the opening of the Aegean Sea has marked the onset of extension in the area

of the Cyclades, with N-S extension rates in the range 4-6 cm yr-1 (Taymaz et al., 1991).

The Aegean is thought to have been in an active extensional setting since at least 33 Ma on

the basis of sediment ages in extensional graben (Gautier and Brun, 1994). Extension

has facilitated the exhumation of high pressure metamorphic rocks, formed during the

Alpine orogeny, along low angle detachment faults (Lister et al., 1984).

0 200km

Levantine Basin

AFRICA

GREECE

TURKEYAegean Sea

TU RK I S H P LATE

ARABI AN P LATE

Black Sea

34º

24º 22º

N

Figure 1-3: Present-day plate configuration for the Eastern Mediterranean showing relative platemotions (adapted from Piper et al., 1996).

At the beginning of the Phanerozoic, the Cyclades were located at the northern

margin of Africa (Figure 1-4), when it was part of the supercontinent Gondwana, and

they have since been subject to two major mountain-building episodes or orogenies. The

first was the mid to late Palaeozoic collision of Gondwana and Laurasia to form the

supercontinent Pangea. The resulting suture zone in the core of Pangea formed the

Variscan mountain chain. After the break-up of Pangea in the Jurassic, the Tethys ocean

developed between the northern margin of Africa and southern Eurasia. This ocean was

destroyed by convergence of the African and Eurasian plates, with final closure occuring

in the late Cretaceous (Smith, 1971; Burchfiel, 1980; Sengör and Yilmaz, 1981) and

subsequent formation of the Alpine mountain chain marking the suture between the

African and Eurasian plates (Seidel et al., 1982). Progressive fragmentation of the

6 Introductionnorthern African plate accompanied this convergence, with slivers of oceanic and

continental crust complicating the shape of the suture zone. Although the relative motions

of the African and Eurasian plates are well-constrained from the Jurassic, based on

palaeomagnetic data and ocean floor magnetic anomaly correlations (Livermore and

Smith, 1984; Dercourt et al., 1986), the motion of the intervening agglomeration of

crustal material that comprises the Turkish plate (and includes the Cyclades) is poorly

constrained. Although detailed palaeomagnetic work is being conducted in this area (see

Morris and Tarling, 1996) the exact nature and position of crustal material is unknown

(cf. Robertson and Dixon, 1984; Robertson et al., 1996). For a recent detailed review of

different tectonic models for the evolution of the eastern Mediterranean the reader is

referred to Robertson et al. (1996). The complex rocks of the Cyclades can thus provide

insights into the operation of both the Variscan and Alpine collisional orogenies, as well

as yielding more general information about the manner in which mountain belts are

created and destroyed.

BALTICA

SIBERIA

IAPETUS OCEAN

GONDWANA

EARLY ORDOVICIAN

LA

UR

ENTIA

TORNQUIST SEA

South Pole

Equator30˚ S

60˚ S

Figure 1-4: Continental reconstruction for the Ordovician from palaeomagnetic data (Torsvik andTrench, 1991), with the position of Laurentia altered, following the determinations of Dalziel (1997).The location of the Cyclades as part of the northern margin of West Gondwana is depicted with darkershading.

Chapter 1 71.3 Regional Geology

The Cyclades form part of the Attic-Cycladic Massif (ACM) (Dürr et al., 1978),

part of a continuous Alpine orogenic belt that can be traced through southern Europe into

the Hellenides and Taurides (Figure 1-5). They form part of the Attic-Cycladic crystalline

belt (Dürr et al., 1978) that extends from Evvia in southeastern mainland Greece to the

Menderes Massif of Turkey (Figure 1-5). The extent of the belt in the region of the

Cyclades is defined by underwater channels (Dürr, 1986) visible in the bathymetric map

of the Aegean Sea (Figure 1-1). The ACM is thought to link the Pelagonian zone of

mainland Greece to the Menderes Massif, Turkey (Dürr et al., 1978; Blake et al., 1981;

Okay, 1984). It comprises a pile of nappes, ophiolites, pelagic sediments and high-P

low-T metamorphic rocks, that were exhumed along low-angle detachments to form

metamorphic core complexes (Lister et al., 1984). A variety of lithologies of different

ages are preserved, the oldest of which have been affected by at least two major

geological events - the Palaeozoic Variscan and Cretaceous-Cainozoic Alpine orogenies.

The geological evolution of this area is difficult to reconstruct because of limited outcrop.

However, due to its median position, reconstruction of the ACM through time is crucial to

our understanding of the geologic relationships between parts of Turkey (Taurides) and

Greece (Hellenides).

CARPATHIANS

APPENINES

DINARIDES

ATLAS HELLENIDES

TAURIDES

PONTIDES

ALPS

CreteSicily Cyprus

BALKANIDE

Africa

TurkeyGreece

Arabia

Eurasia

AC

Figure 1-5: Location of the Attic-Cycladic Massif (AC) in the Aegean Sea, with major tectonic unitslabelled. The extent of the Alpine orogenic belt is shaded (Adapted from Channell and Kozur, 1997).

8 Introduction1.4 Geology of the Cyclades

Although complicated structures make it difficult to correlate between islands, three

main lithological groups can be distinguished (Altherr, 1977; Dürr et al., 1978) as shown

in Figure 1-6 and Figure 1-7.

UPPER UNIT

SERIES

BASEMENT

ophiolite

marblesedimentssediments

granitoid

schist

granitoid

ophiolite

marblesedimentssediments

Mesozoic sediments

not affected by M1 or M2some units affected byCretaceous high-T metamorphism

Mesozoic sediments

affected by M1 and M2

Pre-Mesozoic schists and orthogneisses

affected by M0, M1 and M2

generallyincreasingstructuraldepth

intruded by Miocene granitoids

garnet-mica schists

intruded by Palaeozoic granitoids

granitoid

molasseMiocene molasse

Permo-Triassic limestones

Permo-Triassic limestones

Figure 1-6: Schematic stratigraphic column showing the three main lithological units found in theCyclades. From oldest to youngest these include; the Basement comprised of garnet mica schists intrudedby Palaeozoic granitoids, the Series rocks comprised of Mesozoic sedimentary sequences intruded byMiocene granitoids and the Upper Unit of Mesozoic sediments, ophiolites and granitoids.

1) An “Upper Unit” occurs in the highest structural levels but usually occupies the

lowest topography, and consists of scattered outcrops of fault-juxtaposed medium-low

grade to unmetamorphosed sedimentary and ophiolitic rocks (Roesler, 1978). Fossil

evidence and radiometric dating have yielded Cretaceous ages for sediments and ophiolitic

melanges in Upper Unit rocks of the Cyclades (Dürr et al., 1978; Reinecke et al., 1982;

Patzak et al., 1994) although some marbles as old as Permian have been reported (Marks

and Schuiling (1965) in (Dürr et al., 1978).

2) At structurally intermediate levels, there is a sequence of intensely deformed

Mesozoic platform sediments consisting of a sequence of eclogite-blueschist facies

metamorphosed neritic carbonates, psammitic to pelitic sediments and basic/acid

volcanics, referred to collectively as the “Mesozoic Series” or Series. Only poor age

Chapter 1 9constraints from rare fossil assemblages in these units are available and these range in age

from Triassic to Cretaceous (Dürr et al., 1978). On the island of Syros, part of this unit

is composed of eclogites formed from subducted oceanic crust of unknown age, while on

Naxos, karst bauxites form lenses within the marble units and have been correlated with

Jurassic bauxites found in other parts of the Eastern Mediterranean (Feenstra, 1985).

3) Both the Upper Unit and the Series structurally overlie an inferred pre-Alpidic

“Basement” of schists, gneisses and amphibolites. These are the oldest units found in the

Cyclades, and consist of deformed and metamorphosed granites of presumed Early

Palaeozoic age (Henjes-Kunst and Kreuzer, 1982; Andriessen et al., 1987). The granites

have intruded an older sedimentary sequence, now metamorphosed, of unknown age.

Andros

Tinos

Syros

Kythnos

Serifos

Sifnos

Delos

NaxosParos

Antiparos

Ios

SikinosFolegandros 0 20km

CYCLADES

Mykonos

Upper UnitMiocene GranitesMesozoic Series RocksPre-Alpine Basement

24o30'E

37o30'N

37oN

N

EW

S

Figure 1-7: Simplified geology of the Cyclades showing the distribution of the main tectonic units(adapted from Gautier and Brun, 1994).

The Basement and Series rocks are both intruded by a range of Miocene I- and S-

type granites forming dykes and small plutons on the islands of Naxos, Paros, Tinos,

Mykonos and Serifos (Figure 1-7). These intrusives are thought to be related to

subduction of oceanic lithosphere beneath the Attic-Cycladic complex during the

10 IntroductionOligocene and early Miocene as the Hellenic trench migrated southwestwards to its

present location south of Crete (Altherr et al., 1982; Altherr et al., 1988).

1.5 Metamorphic History of Cyclades

The Pre-Mesozoic basement and the Mesozoic Series rocks have both undergone at

least two regional metamorphic events: a Palaeo-Eocene high pressure blueschist-eclogite

facies metamorphism (M1) followed by an Oligo-Miocene medium pressure greenschist

metamorphism (M2) which has reached amphibolite facies in places (Altherr et al., 1979;

Andriessen et al., 1979; Maluski et al., 1981; Altherr et al., 1982; Henjes-Kunst and

Kreuzer, 1982; Maluski et al., 1987; Wijbrans and McDougall, 1988) (Figure 1-8).

Andros

Tinos

Syros

Kythnos

Serifos

Sifnos

Delos

NaxosParos

Antiparos

Ios

SikinosFolegandros 0 20km

CYCLADES

Mykonos

High P / Low T M1 metamorphismLow / Medium T / Med P M2 metamorphismPartial Melting associated with M2

Miocene GranitesUpper Unit

24o30'E

37o30'N

37oN

N

EW

S

Figure 1-8: Schematic diagram showing the preservation of different grades of metamorphism (modifiedfrom van der Maar and Jansen, 1981; Okrusch and Bröcker, 1990).

M2 reaches its highest grade on Naxos where there is evidence of partial melting

(Jansen and Schuiling, 1976). Both M1 and M2 are related to the Cretaceous-Cainozoic

Alpine Orogeny, but evidence for an earlier M0 event has been described locally in the

Pre-Mesozoic Basement rocks (Henjes-Kunst and Kreuzer, 1982; Franz et al., 1993).

Chapter 1 11This enigmatic amphibolite facies event is poorly constrained, but has been interpreted as

a relict of the Palaeozoic Variscan Orogeny (Henjes-Kunst and Kreuzer, 1982). Local

overprinting caused by contact metamorphism associated with Miocene granitoid

intrusions (M3) is also reported (Altherr et al., 1979; Andriessen et al., 1979; Altherr et

al., 1982; Henjes-Kunst and Kreuzer, 1982). A Cretaceous-aged high temperature

metamorphic overprint is recognised in the Upper Unit of the Cyclades, which has

reached lower amphibolite facies grade (Reinecke et al., 1982; Patzak et al., 1994). This

has not been recognised in the Series or Basement rocks of the Cyclades and so is given

the general tag M? to denote its uncertain place in the metamorphic scheme of the Cyclades

as a whole.

The above is a very simplified summary of the metamorphic history of the

Cyclades, and different islands have experienced different metamorphic grades and

preserve different metamorphic assemblages (see Figure 1-8). There is still debate over

whether there is more than one high-P Alpine metamorphism and more than one

greenschist overprint (Lister and Raouzaios, 1996). The geological and metamorphic

evolution of the Cyclades in relation to a general timescale of geological periods and

events is given in Figure 1-9.

1.6 Structural Development in the Cyclades

The metamorphic history of the Cyclades has been complicated by intense

deformation during the Alpine orogeny and subsequent extension. While there are

differences in the detailed structural histories of each island, a generalised deformational

sequence can be defined, related to the tectonic history of the area (also see Figure 1-9).

D0 is an amphibolite-grade deformational fabric recognised only in Basement units.

D1 is related to the Alpine collision of Africa and Eurasia and has resulted in the

development of at least two sets of isoclinal folds, the first formed during peak-M1

conditions in the lawsonite stability field on Syros (Dixon, 1969) and Ios (Grütter, 1993),

and the second during blueschist metamorphism post-peak M1 (Ridley, 1984).

D2 consists of another period of isoclinal folding related to greenschist facies M2 on

Tinos (Avigad et al., 1988) and mid to upper amphibolite facies M2 on Naxos (Buick and

Holland, 1989; Buick, 1991; Buick, 1991) and is thought to occur in response to

extension of the Aegean crust.

12 Introduction

Figure 1-9: Geologicaltimescale delineating relevanttime periods mentioned in thetext and showing theapproximate timing ofimportant geologic events in theCyclades. The timescale isbased on the AGSO PhanerozoicTimescale (Young and Laurie,1996) with subdivisions oforogenic periods, such as theVariscan, adapted from vanEysinga (1975). The timing ofthe Pan-African orogeny istaken from Black and Liegeois(1993).

PH

AN

ER

OZ

OIC

ARCHAEAN

PALAEOPROTEROZOIC

MESOPROTEROZOIC

NEOPROTEROZOIC

2500

1600

1000

545

PA

LA

EO

ZO

IC

CAMBRIANE

ML

509498490

ORDOVICIAN

434

410SILURIAN

DEVONIAN

E

M

L

384

369

354

CARBONIFEROUS

298

PERMIAN

251270

241230

205

E

L

TRIASSIC

JURASSIC

CRETACEOUS

ME

SO

ZO

IC

E

M

L

E

L

CA

INO

ZO

IC

TER

TIA

RY

PALEOCENE

EOCENE

OLIGOCENE

MIOCENE

PLIOCENE

184

159141

97.5

6554.8

33.723.85.31.78

VARISCAN

CALEDONIAN

LATE ALPINE

EARLY ALPINE

MIDDLE ALPINE

PAN-AFRICAN

M0

M1

M2

BASEMENT

SERIES

UPPER UNIT

Ma

D0

D1

D2 D3 D4

PROT.

Chapter 1 13D3 occurred post-peak M2,, synchronous with Miocene magmatic activity, and

characterised by the development of pervasive mylonitic fabrics associated with

extensional deformation (Buick, 1991). This was partially synchronous with and locally

outlasted by, locally-developed large-scale upright folding developed under lower

amphibolite to greenschist facies metamorphic conditions (Urai et al., 1990; Buick,

1991). It has been suggested that these folds developed in response to doming produced

by removal of overburden during exhumation of the Cycladic core complexes (Grütter,

1993).

D4, locally developed during the late stages of Miocene magmatic activity, consists

of brittle deformation characterised by chloritic brecciation and the development of

pseudotachylites and cataclasites (Jansen and Schuiling, 1976).

The different tectonic slices which comprise the Cyclades have been juxtaposed

along low-angle normal faults, interpreted as detachment faults (Avigad and Garfunkel,

1989; Gautier et al., 1993; John and Howard, 1995). These structures have resulted in

the exhumation of high pressure rocks to form metamorphic core complexes (Lister et al.,

1984; Avigad and Garfunkel, 1989; Urai et al., 1990; Buick, 1991) , defined by a lower

plate comprised typically of Basement and Series rocks, exhumed relative to an upper

plate of Series or Upper Unit rocks.

The Mesozoic metabauxite-bearing marble/schist sequences of the Series rocks

were originally interpreted as material deposited directly on the Basement rocks (Altherr,

1977). However, there is evidence that suggests that the contact between the two units is

tectonic. On Naxos, a tectonic contact is delineated by the presence of ultramafic lenses

thought to define a thrust plane (Jansen, 1973). Katzir (1997) suggests that the ultramafic

rocks are relict peridotites which have been incorporated with underlying rocks prior to

M2 metamorphism during the Eocene (M1) collisional process where an underthrusting

continental slab samples the overriding subcontinental mantle. Jansen and Schuiling

(1976) suggest that the ultramafics may represent remnants of ophiolites emplaced along

pre-metamorphic thrust-planes. A thrust is also thought to separate the basement and

Series rocks on Ios (van der Maar et al., 1981) and Sikinos (Franz et al., 1993), although

on Ios the postulated thrust has been interpreted also as a detachment fault (Lister et al.,

1984), and thus defines a core complex in which the Series rocks represent the upper

plate (rather than the lower plates as identified elsewhere in the Cyclades). Many of the

contacts originally interpreted as thrusts may have been reactivated during Miocene

extension.

It is evident from the preceding discussion that while there are similarities in the

metamorphic and structural histories of the Cycladic islands, their unique histories are

important in unravelling the tectonic history of the region. Hence, the local geology of

each island investigated in this study will be described in the following section.

14 Introduction

1.7 Local Geology

1.7.1 Naxos

Naxos forms a N-S elongate structural dome consisting of all three of the main

Cycladic lithologic units, including a variegated sequence of Mesozoic platform sediments

(Dürr et al., 1978), Variscan basement (Andriessen et al., 1987) and Mio-Pliocene

sediments (Roesler, 1978) (Figure 1-10).

1

2

3

4

56

5

25˚ 35'25˚ 25'

0 1 2 km

Geology of Naxos

N

Granodiorite

Gneiss/Migmatite

Upper Unit

Ultramafics

SchistMarble

Fault

Isograd Surface

+ corundum(420°C)

+ biotite(500°C)

-chloritoid+ staurolite(540°C)

+ sillimanite(620°C)

+ melt phase(670°C)

Zones1 Diaspore2 Chlorite/sericite3 Biotite/chloritoid4 Kyanite5 Sillimanite6 Partial Melt

37˚ 10'

37˚ 0'

Figure 1-10: Geology of Naxos, highlighting the isograd surfaces distinguished by Jansen (1973) andJansen and Schuiling (1976).

Chapter 1 15The Basement and Series rocks of Naxos have been metamorphosed and preserve

the effect of localised high-temperature Barrovian metamorphism (M2) overprinting M1.

The domal structure of the island reveals a series of isograd surfaces which are detailed in

Figure 1-10 (Jansen, 1973; Jansen and Schuiling, 1976), although it should be noted that

some inconsistencies in this pattern have been reported, e.g. for the kyanite isograd

(Buick, 1988). This zonal pattern of increasing metamorphic grade corresponds to

increasing structural depth towards the core. A complete Barrovian sequence of

progressive metamorphism is preserved, with calculated temperatures ranging from 380

˚C in the south-east corner of the island where relict M1 blueschist assemblages are

preserved to ~ 700 ˚C in the core, based on various reaction curves defining the relevant

mineral reactions that produced the isograds (Jansen and Schuiling, 1976).

The rocks consist of alternate layers of marble and pelite with minor intercalations

of basic and ultrabasic material, structurally overlying a gneissic/migmatitic core

dominated by quartzo-feldspathic rocks (Jansen, 1977; Buick, 1988). The marble and

pelite form part of the Mesozoic Series rocks (Dürr et al., 1978) while the core has been

identified as pre-Mesozoic Basement by Andriessen et al. (1987). The proportion of

marble and calc-silicate rocks increases away from the core of the island and titanite has

developed in these units during M2. The metamorphic complex has been ductilely

deformed and intruded during the Miocene by monazite-bearing S-type granites and

zircon/titanite-bearing I-type granitoids both during and subsequent to M2. These

sequences are overlain by a melange of disrupted Mio-Pliocene sediments (Roesler, 1978)

the “Upper Unit”, separated from the metamorphic complex by low angle normal faults

(Lister et al., 1984).

16 Introduction

1.7.2 Paros

An apparent “thermal” dome similar to that of Naxos is also developed on the island

of Paros with M1 gneisses, amphibolites, mica-schists and metabauxite-bearing marbles

undergoing an M2 overprint that ranges in grade from chlorite-sericite to sillimanite

(equivalent to Zones 2-5 on Naxos) (Robert, 1982). The distribution of lithological units

on Paros is illustrated in Figure 1-11.

0 2 km1

Geology of Paros

N

Upper Unit

OrthogneissMarble

Alluvium

Schist/AmphiboliteGranite

37˚ 15'

37˚ 00'

25˚ 15'

Figure 1-11: Geology of Paros (adapted from Papanikolaou, 1980; Robert, 1982)

The Cycladic basement corresponds to the lower group of the Marathi Nappe

defined by Papanikolaou (1980) and is dominated by orthogneisses. The upper unit of

the Marathi Nappe, composed of amphibolites with an inferred Permo-Triassic age and

marbles with an inferred Triassic-Cretaceous age (Papanikolaou, 1980), correspond to

Mesozoic Series rocks. The upper and lower units of the Marathi Nappe are separated by

either a decollement or an overthrust and both have been intruded by Miocene granites.

Chapter 1 17In turn it is overlain by two sequences forming part of the Upper Unit of the Cyclades: a

low grade sequence of marble, phyllites and diabases (the Dryos Nappe); along with non-

metamorphic ophiolites and sediments (the Marmara Nappe).

1.7.3 Ios

The geology of Ios has been well described (van der Maar and Jansen, 1981;

Henjes-Kunst and Kreuzer, 1982; van der Maar and Jansen, 1983; Grütter, 1993) and is

illustrated in Figure 1-12.

0 2 4 km

Geology of Ios

N

OrthogneissGarnet Mica Schist

MarbleSchist

36˚ 45'

36˚ 40'

25˚ 20'25˚ 15'

Figure 1-12: Geology of Ios (adapted from van der Maar and Jansen, 1983)

The lithologies on Ios are broadly grouped as part of either a Pre-Alpine Unit

(PAU) or a Blueschist Unit (BSU) by (Grütter, 1993). The PAU can be correlated with

the Cycladic Basement as defined above and consists of mica schists, metabasites and

granitic orthogneisses, while the BSU contains metapelites, acid gneisses and metabasites

18 Introductionwhich correspond to the Series rocks. The nature of the contact between the PAU and

BSU is unresolved. Henjes-Kunst (1980) - in Grütter (1993), interpreted the contact as

an erosional unconformity with the BSU deposited on top of the PAU, whereas van der

Maar (1983) and Grütter (1993) identified a thrust fault but Lister (1984) interprets the

fault as a detachment. Peak M1 conditions in the BSU are estimated at 9-11 kbar and 350-

400 ˚C on the basis of albite-quartz assemblages with an M2 overprint estimated at 5-7

kbar and 380-420 ˚C from the breakdown of aragonite and sodic pyroxenes (van der

Maar and Jansen, 1983). More recent estimates of P-T conditions using the program

THERMOCALC, distributed with the dataset of Holland and Powell (1990) yield P-T

estimates of 12.6 ± 0.6 kbar at 475 ± 25 ˚C for M1 and ~ 4 kbar and > 400 ˚C for M2

(Grütter, 1993). M1 in the PAU is estimated to be at > 13 kbar and 450-500 ˚C and M2 at

~ 7 kbar and T < 440 ˚C (Henjes-Kunst, 1980 in Grütter (1993)).

1.7.4 Syros

Lithologies on Syros consist of a thick sequence of alternating pelitic schists,

metabasites and marbles (Hecht, 1984) (Figure 1-13). In the northeast of the island, at

structurally highest levels, contain some of the best-preserved M1 blueschist assemblages

in the Cyclades. However, in the south, the lowermost units have been almost entirely

retrogressed to greenschist facies during M2 (Ridley, 1984). Metasediments are mainly

pelitic rather than psammitic and have been interpreted as flysch (Dixon and Ridley,

1987). The metabasite sequences have the appearance of ophiolitic melange (Dixon and

Ridley, 1987) although recent work suggests that the ophiolite sequences can be traced as

stratigraphic units and that their disrupted appearance is due to later deformation

(Ballhaus, pers comm.). The Eocene M1 high-pressure metamorphism is estimated to

have occurred at 450-500 ˚C and > 14 kbar (Ridley and Dixon, 1984; Ridley, 1984). The

geochemistry of the metabasites suggests they were formed in a back-arc setting possibly

at a spreading ridge in proximity to a transform fault (Seck et al., 1996).

Chapter 1 19

0 1 2 km

Geology of Syros

N

MetabasiteOrthogneiss

Marble

Schist37˚ 30'

37˚ 25'

24˚ 55'24˚ 50'

Figure 1-13: Geology of Syros, (adapted from Hecht, 1984).

1.7.5 Sifnos

Like Syros, Sifnos preserves high-pressure metamorphic assemblages in a

sequence of metasediments and metavolcanics. The assemblages are especially well-

preserved in the Eclogite-Blueschist domain (EBD) at the north of the island (Figure 1-

14). Peak metamorphic conditions recorded in this unit are 470 ± 30 ˚C and 15 ± 3 kbar

from meta-acidites containing the critical high-P breakdown assemblage of albite going to

jadeite-quartz (Schliestedt, 1986). This unit is enclosed by two marble units. The

underlying Main Marble Unit contains high-P metamorphic assemblages which have been

variably overprinted by greenschist facies retrogression; this marble is itself underlain by

a metavolcanic-sedimentary unit that crops out in the southern half of the island. This

20 Introductionlower unit has been overprinted by greenschist facies metamorphism and is referred to as

the Greenschist Domain (GSD).

0 1 2 km

Geology of Sifnos

N

Marble

Schist

Alluvium

EBD

GSD

24˚ 40' 24˚ 45'

37˚ 00'

36˚ 55'

Figure 1-14: Geology of Sifnos adapted from Schliested and Matthews (1987).

P-T estimates for M1 in the GSD indicate eclogite-facies conditions at 480-520 ˚C

and 12-15 kbar (Avigad et al., 1992) which is in accord with estimates for the EBD.

Subsequent blueschist and then greenschist overprinting of the unit is thought to have

occurred at ca . 480-500 ˚C and 8-10 kbar, and ~ 450 ˚C and ~ 5.5 kbar respectively

Chapter 1 21(Avigad et al., 1992). Protoliths to both the EBD and GSD are thought to be typical

continental margin sequences, in contrast to the ophiolitic melange preserved on Syros.

The geochemistry of the basic volcanics of the EBD on Sifnos has been investigated by

Mocek (1996) who suggested they have a tholeiitic affinity typical of the early stages of

back-arc basin evolution. The presence of some rocks with bonninitic affinities has been

used to suggest that the rocks developed during the transition from an island-arc to a

spreading centre environment (Mocek, 1996).

1.7.6 Sikinos

Van der Maar et al. (1981) have divided the lithologies of Sikinos into two units, a

lower unit comprised of mica schists that host metaintrusives; and an upper unit of

alternating marbles, schists, metabauxite, basic and ultrabasic rocks (Figure 1-15).

0 1 2 km

Geology of Sikinos

N

Orthogneiss

Marble

Garnet Mica Schist

Schist

25˚ 05' 25˚ 10'

36˚ 40'

Figure 1-15: Geology of Sikinos adapted from Franz et al. (1993).

The lower unit is interpreted as pre-Alpine Basement while the upper unit is

correlated with the Series rocks (van der Maar et al., 1981). Pressure and temperature

conditions during metamorphism have been estimated from the basement rocks as: M0

(pre-Alpine) at 570-650 ˚C and ~ 5 kbar (on the basis of the assemblage staurolite-biotite-

22 Introductionmuscovite-quartz) , M1 (Alpine) at 450-500 ˚C and ~ 10 kbar (using the phengite

geobarometer of Massonne, 1991) and M2 (Miocene) 440-480 ˚C and ~ 5 kbar (using the

plagioclase-hornblende geothermometer of Blundy and Holland, 1990) and the garnet-

hornblende geothermometer of Graham and Powell (1984) (Franz et al., 1993).

1.7.7 Folegandros

Folegandros consists of a core of marble covered at higher structural levels by

schists interlayered with metabasic rocks which are equivalent to Cycladic Series rocks.

The Series units are in contact with fault-bound slices of low grade metamorphosed

sediments (flysch) equivalent to the Upper Unit of the Cyclades (Dürr, 1986) (Figure 1-

16). No basement gneisses are evident on the island.

0 1 2 km

Geology of Folegandros

N

MarbleSchistAlluvium

36˚ 40'

36˚ 35'24˚ 55'24˚ 50'

Figure 1-16: Geology of Folegandros (adapted from Verginis, 1973).

Chapter 1 23

1.7.8 Tinos

Tinos, the third largest island of the Cyclades, is comprised of Series rocks

consisting of a sequence of M1 blueschist metavolcanics and metasediments intercalated

with marble (Figure 1-17) and variably overprinted by greenschist facies retrogression

(M2) on scales of less than one metre (Bröcker, 1990). The Series is intruded by a large

I-type monzogranite and a younger much smaller S-type granite. Low-grade

metasediments correlative with the Cycladic Upper Unit structurally overlie the Series

rocks, and are emplaced as klippen along a low-angle normal fault (Avigad and

Garfunkel, 1989).

0 2 4km

Geology of Tinos

N

Metabasite

Granite

Marble

Schist

Upper Unit

37˚ 30'25˚ 00'

Figure 1-17: Geology of Tinos (adapted from Melidonis, 1980).

To constrain the complex structural and metamorphic evolution of the Cycladic

islands requires a combination of both careful fieldwork and precise geochronology. The

next section describes how geochronology can be applied to reconstruct the Alpine and

pre-Alpine history of the Cyclades.

24 Introduction

1.8 SHRIMP U, Th-Pb Geochronology of Accessory Minerals

Radiometric dating is based on the decay of radioactive isotopes and can be used to

calculate ages from the measured proportions of parent and daughter isotopes (Appendix

D, Section D1). The spontaneous decay of U and Th isotopes to different isotopes of Pb

allows the calculation of ages by several methods. Ages can be determined utilising the

decay of 238U to 206Pb, 235U to 207Pb and 232Th to 208Pb (Section D2). The level of

agreement in the ages determined using these independent decay schemes gives a measure

of the concordance of the systems. Analyses which disagree beyond experimental error

are termed “discordant” and may reflect the open system behaviour of one or more of the

isotopic species being measured. The degree of concordance of a radiometric age can be

assessed by the use of a concordia diagram, with discordance most commonly produced

by Pb loss or U gain in the system (Section D2). One other isotope of Pb exists in nature

which is not produced radiogenically and is termed “common Pb”. Corrections have to

be made for the amount of non-radiogneic lead incorporated into a mineral at its time of

formation (Section D6). Because of the difference in half-life between 238U and 235U, the

ratio 207Pb/206Pb is time-dependent and is also used in age determinations. The relatively

small proportion of 235U which exists now in nature means that the amount of radiogenic207Pb produced by decay of this isotope is also small. This means that age determinations

which rely on the measurement of radiogenic 207Pb, such as the 207Pb/206Pb method, are

inappropriate for use in young samples. For this reason U,Th-Pb dating results reported

in this study will utilise 207Pb/206Pb ages only for those samples older than 1000 Ma,

while the age of younger samples will be quoted as 238U/206Pb or 232Th/208Pb ages.

A variety of analytical techniques may be employed to measure isotopic species, the

most commonly used for U,Th-Pb dating are TIMS (Thermal Ionisation Mass

Spectrometry) and SIMS (Secondary Ion Mass Spectrometry). Advances in spectrometry

have resulted in more precise data being obtained on progressively smaller samples. This

has seen geochronology advance from dating bulk whole-rock samples and mineral

separates, to dating individual minerals, and most recently, to dating zones within

individual mineral grains. Conventional whole-rock isochron methods are still widely

used, e.g., Re-Os, although new methods of precise geochronology are being introduced,

e.g., single grain or ion probe dating methods (Hofmann, 1993). The precise dating of

zones in accessory minerals by the U-Pb method using SHRIMP (Sensitive High

Resolution Ion Micro Probe) can constrain the timing of separate mineral-forming events.

This can provide valuable information on the rate at which tectonic processes (causing

mineral growth) operate. SHRIMP is mainly used to measure the U-Pb systematics of

high-temperature accessory minerals that may record evidence of more than one geologic

event in their internal growth structures and hence provide detailed information

constraining the geological evolution of their host rock. Further details of the SHRIMP

Chapter 1 25technique are given in Appendix D, which describes the design of the SHRIMP

instruments housed in the Research School of Earth Sciences, and discusses the methods

of data collection and data reduction.

U, Th-bearing minerals are not all suitable for geochronology. A number of factors

govern the suitability of minerals for U,Th-Pb dating purposes. Such minerals require

the following characteristics:

1) measurable quantites of radiogenic PbU

2) ability to retain Pb in the crystal lattice (i.e. to resist Pb loss)

3) to incorporate only small amounts of common Pb into their structure so that

radiogenically produced Pb dominates the measured Pb.

4) to occur as relatively common constituents in rocks

5) to have well-constrained formation histories which may be related to geologically

meaningful structures or igneous/metamorphic assemblages

Three accessory minerals which meet these criteria are decribed in the following

sections.

1.8.1 Closure Temperature (Tc)

At this point is necessary to introduce the concept of mineral closure temperature. It

cannot be assumed that isotopic systems remain closed to loss of parent or daughter

isotopes after crystallisation, particularly in metamorphic rocks, where open system

behaviour of isotopes is commonly recorded and can be modelled as a process of volume

diffusion. Dating in many cases thus records an "apparent age" for minerals representing

neither their crystallisation nor complete isotopic closure since closure on cooling is an

exponential process. A mathematical formulation of the transition from open to closed

system behaviour of isotopes in different minerals was developed by (Dodson, 1973)

who introduced the concept of a unique closure temperature (Tc) for minerals.

TE R

ART D a

E T t

T E

A

R T D

a t T t

c

c

=

= =

=

= = =

= = =

ln

.

02

02

0

∂ ∂

∂ ∂

where closure temperature, activation energy,

a value representing the geometry of the mineral,

gas constant, temperature, diffusion coefficient,

characteristic diffusion size, time, cooling rate

Tc is defined as the temperature of a system at a time given by the apparent age of a

mineral, and must be evaluated for a specific cooling rate. Ideally, once the Tc of a

mineral is known, quantitative age estimates for several points along a P-T-t path can be

26 Introductionestablished. In general these age estimates are unlikely to provide information on the

timing of prograde or peak metamorphism, as the mineral will remain above its Tc until it

is on the retrograde path (Cliff, 1993).

A generalised summary of accepted closure temperatures over a range of cooling

rates for the diffusion of Pb from different minerals is given in Table 1-1. Most of the

reported closure temperatures are based on experimental and/or observational data, with

the generally inadequate diffusion data bases implying uncertainties of at least +/- 50 ˚C.

The practice of dating a suite of minerals of known Tc to produce an apparent cooling

pattern is particularly well established for minerals with low closure temperatures (500-

100 ˚C) due to the extensive application of K-Ar, 40Ar-39Ar and fission track studies to

metamorphism. The pattern of cooling from higher temperatures (> 500 ˚C) is less well-

established. An increased use of U-Pb ages should significantly improve the resolution

of the high temperature part of P-T-t paths (Brown, 1993).

Table 1-1: Compilation of Mineral Closure Temperatures for U/Pb isotopes

Mineral Tc Reference

Zircon > 900 °C Lee et al. (1997)

Monazite 675-700 °C

640-730 °C

720-750 °C

> 700-570 °C

Heaman and Parrish (1991)

Mezger et al. (1991) [2°C/Ma cooling]

Copeland et al. (1988) [20°C/Ma]

Spear and Parrish (1996)

Titanite 737 °C

778 °C

Cherniak (1993) [2°C/Ma]

Cherniak (1993) [10°C/Ma]

for grain with 0.5 cm radii

1.8.2 Application of Geochronology to Metamorphic Processes

Establishing the chronology of metamorphic events places important constraints on

the thermal evolution of a terrane and the age of its protoliths, and also provides

information vital for the construction of accurate P-T-t paths. With the recent advent of

high precision dating techniques capable of analysing a range of metamorphic minerals,

accurate and reliable time constraints can be placed on metamorphic processes (e.g.,

Mezger, 1990) A range of minerals with different Tcs can be dated, and the time

constraints used to construct pressure-temperature-time (P-T-t) paths. These paths trace

the tectonic evolution of crustal segments by relating mineral Tc to particular geological

processes or events. Hence, the derivation of geologically meaningful P-T-t histories

requires careful integration of radiometric ages with evidence from other disciplines..

Precise U-Pb dating of accessory minerals, with their generally higher Tc than minerals

Chapter 1 27dated by other systems, may constrain the high temperature portion of P-T-t paths and in

some cases constrain the prograde metamorphic path (Smith and Barreiro, 1990; Dirks

and Hand, 1991). The interpretation of geochronological results however, relies on

careful selection and characterisation of samples and also on assumptions about the

physical and chemical properties of the minerals dated. The three U-bearing minerals

investigated in this study are zircon, monazite and titanite (for analytical procedures refer

to Appendix D). Each mineral has its own distinctive characteristics and Tc with the

minimum temperature at which Pb is lost by diffusion from the crystal lattice decreasing

in the order zircon > monazite > titanite (see Section 1.8.1). For this reason zircon has

often been used to constrain the timing of “peak” metamorphic conditions (e.g., Kröner

and Jaeckel, 1995), although the mechanisms of metamorphic zircon growth remain

poorly understood (e.g., Fraser et al., 1997; Roberts and Finger, 1997) and the

possibility that metamorphic minerals grow at temperatures below their Tc can further

complicate age interpretation.

Accessory phases may develop in response to metamorphic conditions during:

1) subsolidus metamorphic breakdown reactions, e.g., monazite in pelites (Smith and

Barreiro, 1990); titanite in calc-silcates (Hunt and Kerrick, 1977);

2) fluid infiltration (metasomatism) in permeable rock units, e.g., titanite in calc-silicates

(Cliff, 1993);

3) anatexis, e.g., partial melt produced zircon (Roberts and Finger, 1997) or monazite

(Watt and Harley, 1993);

4) deformation, e.g., growth of secondary titanite in shear zones (Johansson and

Johansson, 1993) or precipitation of zircon (Wayne and Sinha, 1992). This process may

be considered to be another version of fluid infiltration, in this case focussed along and

adjacent to shear zones where permeability has been deformation- or reaction-enhanced

(Rumble and Spear, 1983).

Fluid flow transfers heat into the crust, causing extensive metasomatism which

affects crustal rheology. Fluid infiltration can occur under a variety of conditions and

may be focussed preferentially into certain rock units depending on their permeability and

chemistry [Brocker, 1990 #1573]. In particular, the influence of fluid flow and the

growth of accessory minerals as a result of hydrothermal processes can complicate the

construction of P-T-t paths in metamorphic terranes. Minerals that developed in response

to ambient conditions such that isotopic ages reflect the time that these minerals passed

through their closure temperature can easily be integrated into a P-T-t scheme. If,

however, minerals are produced by interaction with an externally-derived fluid, then their

conditions of formation are more difficult to constrain. Fluid flow may occur in response

to identifiable stages of the metamorphic process, for example, during the development of

28 Introductiona contact aureole around an intrusion or by channellisation along a shear zone.

Determining the temporal and spatial characteristics of fluid flow can yield information on

segments of the metamorphic P-T-t-fluid composition (Xi) paths and hence on the thermal

and deformational histories of orogenic belts. Recognition of the relationship between

fluid infiltration and metamorphism is critical if meaningful ages are to be obtained and

integrated into the tectonic framework of an area (cf. Williams, 1996). The identification

of fluid flow in the metamorphic environment is described in Section 1.12.

The next section details the characteristics of the U-Pb bearing accessory minerals

analysed in this study and their geological significance. Each mineral, zircon, titanite and

monazite, shows a characteristic style of response to different ambient conditions. The

mechanisms by which U-bearing accessory minerals such as zircon, monazite and titanite

may form, and the ages derived from such minerals, are also described. The term

“metamorphic”, when applied to mineral growth, is a general term which is used to

encompass any zircon, monazite or titanite produced during the operation of metamorphic

processes, regardless of the mechanism involved.

1.9 Zircon

Zircon (ZrSiO4) has a tetragonal structure, with Zr4+ in eight-fold co-ordination, and

Si4+ in four-fold co-ordination (Figure 1-18). U is accepted into the zircon structure due

to its similarity to Zr in ionic charge and radius while Pb is largely excluded. Zircon has

long been the favoured mineral for precise U-Pb age determination because it is very

stable, surviving weathering, erosion, sedimentation and high grade metamorphism

(Jäger, 1979). The refractory behaviour of zircon during metamorphism is well

documented by the presence of inherited radiogenic Pb components (Bossart et al.,

1986). Zircon is remarkably chemically stable, which prevents complete erasure of

primary Pb even during crustal anatexis (Pidgeon and Aftalion, 1978). Complete

resetting of the U-Th-Pb system in zircon has rarely been suggested and has not been

documented (Silver, 1991). Although Pb loss may occur during thermal disturbance of

zircons, the timing of this episode of Pb loss can be calculated and may be of tectonic

significance. Due to the slow rates of diffusion of Pb from the zircon structure, Pb loss

often results by fast pathway diffusion of Pb along fractures produced by radiation

damage in old, or high U zircon grains. Zircon contains only minor amounts of common

Pb. This means precise ages for zircon can be calculated even if the initial Pb isotopic

composition is not well known (Cliff, 1985). One complication to U-Pb dating of zircon

is the recognition of the existence of unsupported radiogenic Pb in some zircons,

unrelated to either U loss or Pb gain (Williams et al., 1984).

Chapter 1 29

Figure 1-18: Zircon structure shown by perspective polyhedral and ball and stick representation withruled tetrahedra equivalent to SiO4 groups and shaded polygons representing ZrO8 groups. Uraniumsubstitutes into the Zr position (adapted from Speer, 1980).

Interpretation of ages derived from zircon in the metamorphic environment is

complicated by our imprecise knowledge of the nature of Pb diffusion in zircon, and the

mechanisms by which zircon is formed. The mineral reactions which cause zircon

growth during metamorphism are poorly constrained (Fraser et al., 1997; Pan, 1997) and

difficult to separate from other processes such as crystallisation from partial melts,

precipitation from hydrothermal fluids or the complicating effects of solid-state

recrystallisation. Each of these mechanisms can be broadly termed “metamorphic” but are

related to the process of metamorphism in different ways. “Metamorphic” zircon can be

recognised by its morphology, producing generally well-facetted euhedral grains or

overgrowths on pre-existing grains with shorter aspect ratios than igneous zircons, and

without the surface pitting indicative of abrasion that is common to sedimentary zircons.

Because the pressures and temperatures associated with new zircon growth during

metamorphism are often unconstrained, zircon ages can not easily be used to construct P-

T-t paths and the interpretation of U-Pb zircon ages often need to be checked against age

data from other minerals and isotopic systems (See Section 1.9.1).

The refractory nature of zircon and the presence of inherited cores in zircon grains

also complicates dating, especially when within-grain analysis techniques are not

employed, resulting in geologically meaningless mixed ages being produced (Williams,

1992). The zonation visible in most zircon grains reflects boundaries between zones of

different geochemical composition that precipitated at different times. The morphology of

30 Introductionthese zones can provide information concerning the multistage history of the grain, and

allow sedimentary, igneous and metamorphic zircon to be distinguished (Hanchar, 1995).

Interpretation of these zones is complicated by the fact that sometimes no new zircon

growth will occur during a metamorphic episode (Cliff, 1985) so a complete tectonic

history can seldom be defined.

1.9.1 Metamorphic Zircon

Metamorphic zircon can be in the form of entirely newly crystallised zircons or as

zircon overgrowths or outgrowths nucleated on pre-existing grains (Saxena, 1966).

Morphological studies of detrital zircons which have undergone high grade metamorphic

contact or regional metamorphic conditions show that zircon becomes progressively

refacetted and euhedral at high temperatures with obvious new growth occurring on pre-

existing cores (Poldervaart, 1955; Gastil et al., 1967; Davis et al., 1968). It can usually

be distinguished as metamorphic growth by its morphology, since metamorphic zircons

have well-developed prismatic faces compared to detrital grains and are often complexly

facetted with smaller aspect ratios than typical igneous zircons. In cases where

metamorphic zircon growth has been reported, it is almost invariably unzoned or displays

only broad zonation patterns (Claoue-Long et al., 1990; Dirks and Hand, 1991; Wayne

and Sinha, 1992; Williams, 1996; Pan, 1997) quite distinct from the oscillatory zoning

patterns found in zircon grown by magmatic precipitation (Pidgeon, 1992). The

exception to this rule would be zircon grown during metamorphism via crystallisation

from partial melts, which is essentially igneous zircon (Roberts and Finger, 1997)

although unzoned zircon from granulite-grade migmatites has been noted (Buick, pers

comm.). In cathodoluminescence, metamorphic zircons are generally unzoned, with low

luminescence often associated with their relatively high uranium contents (this study).

There are some indications of distinct differences in the chemistry of metamorphic zircons

in comparison to magmatic zircons, with metamorphic zircons showing increased Hf and

reduced Y (Vavra et al., 1996; Pan, 1997) and typically low Th/U ratios (Williams and

Claesson, 1987).

The formation of metamorphic zircon was once considered to occur only during

very high grades of metamorphism (Gastil et al., 1967; Davis et al., 1968; Marshall,

1969; Vavra et al., 1996). It has been observed that the morphology of zircons does not

change under amphibolite-facies conditions (e.g., Vavra and Hansen, 1991), and that

metamorphic zircon is not found in all zircon-bearing lithologies even at the same grades

of metamorphism (this study Eckelmann and Poldervaart, 1957; Marshall, 1969; Vavra et

al., 1996). It follows that factors other than metamorphic grade are important in the

production of metamorphic zircon. The formation of metamorphic zircon has been

attributed to a range of different processes including: (1) crystallisation from melts

associated with high grade metamorphism; (2) breakdown of pre-existing radiation-

Chapter 1 31damaged (metamict) zircons to form new zircon (Marshall, 1988); (3) breakdown of

mineral phases to release zirconium and silicon according to well-constrained reaction

processes (Vavra et al., 1996; Fraser et al., 1997; Pan, 1997); (4) fluid infiltration and

hydrothermal activity fluid infiltration (e.g., Mumpton and Roy, 1961; Marshall, 1969;

Claoue-Long et al., 1990; Claoue-Long et al., 1992; Kerrich and King, 1993; Kerrich

and Kyser, 1994; Yeats et al., 1996); or (5) solid-state recrystallisation/replacement

processes (Davis et al., 1968; Marshall, 1969; Black et al., 1986; Pidgeon (1992);

Chiarenzelli and McLelland, (1993); Black and Hoskin, in prep.). Each of these

mechanisms can be related to the process of metamorphism in different ways. Regardless

of the process responsible, metamorphic zircon growth is only rarely reported in rocks

which have not experienced granulite-grade metamorphism.

In the following sections, mechanisms of producing zircon that will record the

timing of some aspect of the metamorphic process are described.

1.9.1.1 Melt Crystallisation

Zircon may crystallise from melts associated with high grade metamorphism, a

process which involves the initial resorption of existing zircon during partial melting

(Davis et al., 1968; Watson and Harrison, 1983), breakdown of Zr-bearing mineral

phases and subsequent growth of new zircon grains or zircon overgrowths nucleated on

pre-existing grains during melt crystallisation. In this situation the zircon produced will

record some stage of the partial melting process and morphologically and chemically the

zircons will appear the same as igneous grains (although note the observation of Buick

(pers comm) of unzoned zircon overgrowths in a migmatite). Factors affecting the

mobility of Zr have an influence on the timing of crystallisation of magmatic Zr-bearing

phases such as zircon. Experimental studies by Dietrich (1968) showed that addition of

sodium in the form Na2SiO5 or NaF to artificial magmas kept Zr in liquid solution. This

relegates zircon to a late-stage crystallisation product associated with the final stages of

magmatism, suggesting that zircon formed in metamorphic rocks in response to partial

melting processes may be a late stage product recording the post-peak timing of

metamorphic activity (cf. Roberts and Finger, 1997). Other experimental studies have

also shown that Zr is progressively enriched in differentiating alkalic and F-rich magmas

(e.g. Watson 1979). It has also been shown that high water contents promote zircon

formation (Marshall, 1969) suggesting that partial melting under water-saturated

conditions, such as those experienced on Naxos (Buick, 1988), would favour new zircon

growth.

1.9.1.2 Net-transfer reactions

Zircon growth due to the breakdown of Zr and Si-bearing minerals involves solid-

state nucleation and crystallisation of new zircon in close proximity to the reaction site,

32 Introductionand, if the controlling net-transfer reaction can be identified it should be possible to

constrain the P-T conditions of metamorphic zircon growth (Vavra et al., 1996; Fraser et

al., 1997; Pan, 1997). This process requires not only the breakdown of Zr and Si-bearing

minerals but also a mechanism to trigger the growth of new zircon. Pre-existing zircon

can breakdown to release zirconium and form zircon and xenotime according to the

reaction (Pan, 1997):

Zr,Hf,Y,REE Si,P O Zr,Hf SiO Y, REE PO

igneous zircon metamorphic zircon xenotime4 4 4( )( ) → ( ) + ( )

Other minerals which contain zirconium as an essential constituent and could

conceivably breakdown to release zirconium are listed in Aja et al., (1995). The more

common zirconium-bearing accessory minerals are baddeleyite and zirconolite, which are

found mainly in silica deficient rocks. Davidson and van Breeman (1988) and Black et al.

(1991) have reported zircon replacement of baddeleyite (ZrO2) during alteration of mafic

magmas, while Ferry (1996) delineated zircon-in isograds related to baddeleyite

breakdown during contact metamorphism on Skye. Baddeleyite and zirconolite will

breakdown to form zircon (Pan, 1997) according to the equations:

ZrO Si 2O ZrSiO

baddeleyite zircon

24 2

4+ + ( ) →+ −

CaZrTi O 2Si 4O ZrSiO CaTiSiO TiO

zirconolite zircon titanite rutile2 7

4 24 5 2+ + → + ++ −

Some major minerals contain sufficient zirconium to supply Zr for metamorphic

zircon growth. In granulite-grade rocks zirconium may be preferentially incorporated into

orthopyroxene (1000-2000 ppm Zr) (Marshall, 1969), along with high levels in

coexisting clinopyroxene (200-500 ppm) and hornblende (200 ppm). Black et al. (1991)

suggested that clinopyroxene could be a Zr donor in mafic rocks whereas Rubin, et al.,

(1993), considered that the sodic pyroxene, aegirine, and the amphibole arfvedsonite, are

potentially significant donors, showing that these minerals contain up to 3.2 wt% and

0.95 wt% Zr respectively. A recent study by Fraser et al. (1997) has revealed that garnet

and hornblende from some granulite-facies rocks have relatively high Zr abundances up to

130 ppm for hornblende and 55 ppm for garnet.

1.9.1.3 Hydrothermal Activity

Hydrothermal activity could include processes such as:

a) fractionation of zirconium from crystallising melts into a residual fluid phase

which can be transported away from the reaction site to produce new metamorphic zircon

growth (Dirks and Hand, 1991).

b) deformation in a shear zone (Wayne and Sinha, 1988) with associated fluid flow

focussing (Dirks and Hand, 1991).

Chapter 1 33Unlike the granulite facies, which is characterised by very low water contents

(Yoder and Tilley, 1962), amphibolite facies metamorphic conditions are generally

associated with a high water content which favours zircon crystallisation (Marshall,

1969). This would enhance zirconium mobility promoting the formation of hydrated

complexes such as (Saxena, 1966):

ZrSiO H O ZrO(OH) + SiO4 2 2 2+ ⇔

Zirconium mobility associated with hydrothermal processes relies on the availability

of zirconium and silica (depend on the breakdown of pre-existing Zr and silica-bearing

mineral phases) and release of these into a fluid to be transported away from reaction site.

It also relies on a mechanism for precipitation.

1.9.1.4 Solid-state Recrystallisation

Recrystallisation of zircon has been described as the refaceting of pre-existing

grains without the addition of new material (Marshall, 1969). Davis et al. (1968)

suggested that recrystallisation of existing zircon would preferentially occur along grain

boundaries. Pidgeon (1992) described this as a replacement texture which resets U-Pb

isotope systematics and changes other geochemical features of the zircon. In particular,

recrystallised zircon seems to show a depletion in LREE compared to the original zircon

grains (Black and Hoskin, in prep.). Distortion of the zircon lattice due to incorporation

of trace elements (Koppel and Sommerauer, 1974) will favour the recrystallisation

process (Pidgeon, 1992) as it makes the lattice unstable. Chiarenzelli and McClelland

(1993) suggested a direct association between temperature and zircon recrystallisation

where it occurs in rocks found at T > 750˚C, but rocks with the same petrography and

geochemistry found at T < 750˚C lack recrystallised zircon. The degree of

recrystallisation is related in an, as yet unquantified way to factors such as pressure,

temperature, textural relationships, fluid availability and time (Pidgeon, 1992). The

potential effects of deformation could also important. The significance of zircon

recrystallisation in response to metamorphic processes is difficult to assess unless the

term “recrystallisation” is properly defined (for a discussion of different types of

recrystallisation see Drury and Urai, 1990). The zircon recrystallisation process most

workers describe is one of replacement rather than grain boundary migration. This is an

important distinction because replacement growth may occur in response to fluid

infiltration, which may or may not be deformation-dependent, while grain boundary

migration is driven by defects in direct response to deformation.

1.9.2 Th/U chemistry of zircons

Low Th/U ratios are considered to be characteristic of metamorphic zircon

(Williams and Claesson, 1987), which suggests that thorium is preferentially excluded

from the zircon lattice relative to uranium. This can be explained in terms of the differing

34 Introductionionic radii of these cations that both substitute for Zr4+ in the zircon lattice. The relevant

ionic radii are Zr4+ <U4+ <Th4+ = 0.84 <1.01 < 1.05 Å (Berger, 1991). It has been

recognised that due to its larger ionic radius Th4+ is depleted in zircon relative to U4+ in all

zircons relative to the abundance of these elements in the crust (Ahrens et al., 1967;

Mattinson, 1973). The even stronger preferential exclusion of Th4+ from the lattice of

metamorphic zircons suggests that there is some other control over cation substitutions or

that there is little Th available for incorporation into the zircon lattice. While scavenging

of Th by another phase common in the metamorphic environment might cause depleted

Th/U ratios in zircons, it should be noted that low Th/U ratios have not been reported in

zircons from granites that contain monazite (a potential Th-scavenger). As noted in

Chapters 4 and 5, this generalisation does not apply to zircons produced by hydrothermal

activity which show similar Th/U ratios to igneous zircons (Claoue-Long et al., 1990;

Black et al., 1991; Yeats et al., 1996). The Th/U ratios for zircons from all samples

analysed during this study were combined and are illustrated in Figure 1-19.

0 20 40 60 80 1000

1.0

0.8

0.6

0.4

0.2

Age (Ma)

Th/U

0 100 200 300 400 5000

0.5

1.0

1.5

2.0

Figure 1-19: Th/U ratios from all zircons younger than 100 Ma with shaded area illustrating the Th/Uratio from magmatic rocks. Inset shows Th/U ratios for all zircons younger than 500Ma.

Chapter 1 35This figure shows the marked decrease in Th/U ratios for zircons aged between 70

and 20 Ma, a time corresponding to the onset of Alpine collisional orogenesis in the

Aegean region. Ages older than 70 Ma have variable Th/U ratios, although, as noted in

Chapter 5, characteristically low Th/U ratios can not be used in isolation to identify zircon

produced by metamorphic processes. Ages younger than 18 Ma are mainly from Miocene

intrusive rocks which show a range in Th/U ratios characteristic of igneous zircons

(Ahrens et al., 1967; Heaman, 1990). The outliers include five data points between 16 to

70 Ma from samples SK9603 and 89642. Both are metabasic rocks. SK9603 is a

metabasite from Sikinos, described in Chapter 4, while 89642 is the Syros ophiolite

described in Chapter 5. As there is no relationship between U content and age in these

samples, Pb loss is not thought to be responsible for producing these ages and they may

reflect metamorphic zircon growth although this is not apparent from the morphology of

these grains. It has been observed that metamorphic zircons from high-grade mafic

samples often have variable Th/U ratios, even when associated zircon growth in felsic

samples exhibits very low Th/U values (Buick, pers comm.).

1.10 Monazite

Monazite ((Ce,LREE,Th,U)PO4) is a LREE-enriched phosphate that commonly

occurs as an accessory mineral in peraluminous-metaluminous granitoids, in low and high

grade Ca-poor pelitic schists and gneisses, as a rare accessory in sedimentary rocks and

as a detrital mineral in beach sands (Parrish, 1990). Monazite is monoclinic and consists

of distorted PO4 tetrahedra with either Ce, LREE, Th or U equidistant from nine oxygen

atoms (Deer et al., 1992) (Figure 1-20). Monazite lends itself to U, Th-Pb dating as it

has favourable U/Pb and Th/Pb ratios with generally high U and Th contents (often in

excess of zircon) and low common Pb concentrations making it relatively insensitive to

corrections for different common Pb compositions (Childe et al., 1993). This makes it

particularly suitable for dating young events (< 30 Ma) as it can accumulate a measurable

quantity of Pb in only a few million years (Smith et al., 1992). It also has the advantage,

like other phosphates, of not accumulating significant radiation damage as silicate

minerals (like zircon) commonly do (Silver, 1990).

As monazite favours the incorporation of Th relative to U, Th/Pb ages can also be

used for dating purposes. However the disadvantage of high Th contents in young

monazite is the presence of considerable excess 206Pb from 230Th disequilibrium

(Scharer, 1984) (see Appendix D). This problem can be largely overcome by the use of

Th-Pb rather than U-Pb ages (e.g., Barth et al., 1989). Inherited monazite has recently

been identified (Copeland et al., 1988) and appears to be a relatively common feature in

igneous environments (Parrish, 1990). However inheritance in monazite is not as

common as in zircon, and, when used with caution, monazite can often be used to attain

36 Introductionthe crystallisation age of granites where inheritance or Pb-loss is a significant problem in

zircon (Williams et al., 1983).

P

o

P

o

REE

Figure 1-20: Monazite structure represented by perspective polyhedra showing chains of REE polyhedra(white) and chains of distorted PO4 tetrahedra (shaded). The ball and stick figure shows the compositionand linking of P and REE with O atoms (adapted from Ni et al., 1995).

It has been suggested that U-Pb monazite ages can record the timing of prograde

metamorphism (Smith and Barreiro, 1990). This suggestion results from the observation

that monazite may grow below its closure temperature (see Section 1.8.1) and so U-Pb

ages will record the growth of monazite during prograde metamorphism (provided peak

metamorphism does not exceed the Tc). This property of monazite seems to be associated

with its instability at moderately low grades of metamorphism. It has often been observed

that while detrital monazite can be identified in sedimentary rocks, once biotite grade

metamorphism (greenschist facies) is reached monazite disappears only to reappear as a

metamorphic mineral at staurolite-grade (525 ± 25 ˚C) (low-medium amphibolite facies)

(Smith and Barreiro, 1990). Monazite then appears to remain stable (resistant to isotopic

resetting), even above its inferred closure temperature, into sillimanite grade

metamorphism (Smith et al., 1992). Harrison et al. (1994) have estimated that the

monazite "isograd" occurs at 450-500 ˚C, and below these temperatures monazite is

thought to dissolve, with LREE and phosphate being redistributed to form allanite,

xenotime and apatite.

1.10.1 Metamorphic Monazite

There are several ways in which monazite can be produced via metamorphic

processes (Pan, 1997), and one possible monazite-producing reaction includes the

breakdown of pre-existing titanite according to the equation:

Chapter 1 37

titanite + CO PO + OH

monazite + quartz + calcite + rutile + epidote + bastnasite + chlorite3

-( ) + ( ) ( ) →− −2

4

3

Pan (1997) also suggested that the common occurrence of monazite as clusters of

anhedral grains in high-grade gneiss indicates that monazite can form from pre-existing

REE-rich minerals such as allanite (Smith and Barreiro, 1990; Bingen et al., 1996) or

apatite (Wolf and London, 1995).

Monazite can be produced by the breakdown of co-existing allanite, hornblende and

titanite at conditions close to the clinopyroxene-in isograd in granulite-facies rocks

(Bingen et al., 1996) according to the following balance of REE contents:

3 HREE O 3 LREE O 2Ca PO F,OH 6SiO

6 LREE PO 2Ca HREE SiO F,OH 6CaO

2 3 2 3 5 4 3 2

4 2 3 4 3

( ) + ( ) + ( ) ( ) + ⇔

( ) + ( ) ( ) ( ) +

where hornblende and titanite provide HREE that are taken in by apatite, and

allanite provides the LREE that are taken up by monazite. Bingen et al. (1996) suggest

that evidence of co-existing relict titanite and metamorphic monazite in upper amphibolite

facies orthogneisses suggests monazite was produced from breakdown of allanite and

titanite from the pre-existing granite protolith.

In melts, the difference in relative solubility between apatite and monazite will result

in monazite precipitation from dissolution of apatite according to the equation (Wolf and

London, 1995):

apatite liquid monazite liquid+ → +

Cheralite-rich high-Th monazite-(Ce) may be produced in late stage residual melts.

It requires the charge-balanced substitution of Th into the monazite structure according to

(Watt, 1995):

Th Si REE P

Th Ca 2REE

4 4 3 5

4 2 3

+ + + +

+ + +

+ ⇔ +

+ ⇔

Watt (1995) first reported Th-rich monazite as a metamorphic phase in granulite-

facies migmatites, where it had previously only been recorded in granitic pegmatites (e.g.,

Mannucci et al., 1986; Wark, 1993). The existence of cheralite-rich monazite in granulite

facies rocks has been related to the influx of late-stage residual fluids (Fitzsimons, 1996).

Hydrothermal processes may produce metamorphic monazite by the exsolution of

fluorapatite according to the following equation (Pan, 1997);

Ca,REE P,Si O F Ca 2P 6O

Ca,REE P,Si O F REEPO SiO12

2 5 2

12 4 2

( )( ) + + →

( )( ) + +

+ + −

38 IntroductionThe above discussion indicates that there are many ways of producing monazite

associated with metamorphic processes and it is important to distinguish how the

monazite formed.

1.11 Titanite

Titanite (CaTi[SiO4](O,OH,F)) is a CaTi-bearing silicate mineral that occurs as an

accessory phase in a range of igneous rocks and which may develop during

metamorphism in marbles, Fe-Mg rich schists, gneisses and impure calc-silicates (Deer et

al., 1992). It is also, rarely, found as detrital grains in sedimentary rocks. Titanite is

monoclinic with main structural units consisting of chains of corner-sharing TiO6

octahedra linked by SiO4 tetrahedra which share the remaining oxygen atoms (Figure 1-

21). This framework encloses Ca in irregular seven coordinated polyhedra which share

edges and corners (Deer et al., 1992).

Ca

Ti

Ti

O

Figure 1-21: Perspective polyhedral and ball and stick representation of the titanite structure, showinglinked chains of Ti-octahedra connected by isolated Si-tetrahedra. Ca is located in a seven-coordinated sitebetween the chains and is the main site for substitution of uranium, thorium and lead. The ball and stickrepresentation illustrates the structure of the Ti octahedron (adapted from Ribbe, 1980).

Chapter 1 39Titanite has favourable Pb/U ratios with U (and also Th) substituting into the site

normally occupied by Ca (Ribbe, 1980). Titanite generally contains more common Pb

than either zircon or monazite as this may also substitute into the Ca site (Ribbe,

1980).Titanite often appears to behave more concordantly than zircon (Tilton and

Grünenfelder, 1968) especially at low temperatures (Krogh and Keppie, 1990). One of

the advantages of using titanite for U-Pb dating is its widespread occurrence in both

metamorphic and igneous rocks. Titanite is the most prevalent radioactive mineral in

common plutonic rocks (Silver, 1990) and, unlike zircon and monazite, titanite

crystallisation in metamorphic rocks can be related to well known chemical reactions that

can be directly linked to the metamorphic reaction history of the rock (Cliff, 1993).

Titanite recrystallisation can also be related to deformational structures (Johansson and

Johansson, 1993). U-Pb systematics in titanite remain closed during low-medium grade

metamorphism unless titanite is directly involved in metamorphic reactions, in which case

it records the timing of the metamorphism. Titanite is increasingly being dated by

conventional U-Pb TIMS to give apparent cooling ages for igneous rocks (Hanson et al.,

1971) and to date medium-high grade metamorphic terranes (Rawnsley, 1987; Parrish,

1989). Titanite ages are used to constrain the temperature interval 500-670 ˚C - (Mezger

et al., 1991) on P-T-t paths (see section 1.8.1). This information can be used in turn, to

suggest cooling rates and constrain tectonic uplift histories. However Getty and Gromet

(1990) have cautioned against the use of titanite for constraining cooling curves,

suggesting that titanite might record partial resetting (a thermal disturbance) rather than

cooling through its closure temperature.

More than one generation of titanite may be present in a rock enabling dating of a

number of tectonothermal episodes (e.g., Getty and Gromet, 1990) Grapes and

Watanabe (1992) report the occurrence of both detrital and authigenic titanite in a

metagreywacke from the Southern Alps, New Zealand. The authigenic titanite has

replaced detrital titanite and ilmenite during metamorphism of the rock, so that both an

apparent provenance age (thought to record an early high P/T event) and a later

metamorphic age could be obtained from the sample. The existence of inheritance in

titanite was predicted (e.g., Cliff, 1993) and has recently been cited (Corfu et al., 1994;

Zhang and Sharer, 1996). Within-grain SHRIMP U-Pb analyses of titanite may enable

dating of retrograde metamorphism eg. overgrowths on titanite may form by release of Ti

during retrogression of minerals like biotite and hornblende at different temperatures

(Cliff, 1993).

1.11.1 Metamorphic Titanite

Titanite is a common accessory mineral in metamorphic rocks, particularly in calc-

silcates, and may form from the well-established reaction (Hunt and Kerrick, 1977):

40 Introduction

calcite + rutile + quartz ⇔ titanite + CO2

Rutile is a common accessory mineral in the high-P mineral assemblage rutile-

calcite-quartz (Taylor and Coleman, 1968) and titanite generally forms as a result of high

temperature overprinting of rutile-bearing mineral assemblages. Titanite may also be

precipitated from Ti-bearing fluids, especially where these react with calcite-rich rocks

(Cliff, 1991).

1.12 Stable Isotopes

It is clear from the preceding discussion that fluid infiltration and hydrothermal

activity can play an important role in the development of U,Th-bearing accessory

minerals. The level of involvement of fluids in medium to high grade metamorphism has

been the subject of considerable debate. The pathways and mechanisms of fluid flow are

not well-understood, but evidence of fluid flow can be found in the development of new

hydrous mineral assemblages, the homogenisation of chemical and isotopic signatures of

pre-existing lithologies and scale invariant mass transfer (e.g., Etheridge and Cooper,

1981; Cartwright and Buick, 1995). Stable isotopes are often used to provide evidence of

regional-scale fluid infiltration in terranes such as Naxos with large 18O depletions relative

to ordinary sedimentary stable isotope ratios (Rye et al., 1976; Baker and Matthews,

1994; Baker and Matthews, 1995).

Stable isotope geochemistry relies on the fact that different isotopes have different

physical and chemical responses to geological processes. These differences arise due to

quantum mechanical effects related to the masses of the different isotopes (Hoefs, 1987).

During a chemical reaction, molecules containing a higher proportion of lighter isotopes

will react more readily than molecules with more of a heavier isotope. Kinetic processes

and isotope exchange reactions may produce isotope fractionation where isotopes are

partitioned preferentially between two substances with different isotope ratios. For

oxygen isotopes the 18O/16O ratio varies in nature by about 100% while carbon isotopes

(13C/12C) also show large variations due to fractionation. The characteristic δ18O and δ13C

values of different fluids and lithologies is shown in Figure 1-22. The relative differences

in isotope abundances in different substances are measured as relative to an arbitrary

international standard, where δ values are calculated by:

18

16

18

16

18

16

1000

O

O

O

O

O

O

initial standard

standard

−

×

Chapter 1 41

40 30 20 10 0 -10 -20 -30 -40 -50 -60 -70

δ18Ο in %

Ocean Water

Sedimentary Rocks

Metamorphic rocks

Granitic rocks

Basaltic rocks

Meteoric Waters

δ13 C in %

40 30 20 10 0 -10 -20 -30 -40 -50 -60 -70

Carbonatites

Air CO2

Marine carbonates

Freshwater carbonates

Marine + non-marine organisms

Sedimentary/organic material

Figure 1-22: Typical δ18O and δ13C values for different rock-types and water (from Hoefs, 1987).

Metamorphic reactions in rocks often occur in the presence of a fluid phase, usually

dominated by water and carbon dioxide in varying proportions associated with reactions

involving carbonates and OH-bearing minerals. The effect of fluid-rock interaction is to

shift oxygen and carbon isotope ratios of both fluid and rock away from their initial

values. If large volumes of fluid are involved, the isotopic composition of the fluid

reservoir will not alter significantly whereas that of the rock will shift appreciably. The

large differences often found in the initial isotope signatures of fluid and rock enable the

levels of isotopic equilibration to be determined and so the influence of externally-derived

fluids during metamorphism can be identified. The application of stable isotope studies in

regional metamorphic terranes is often used to constrain the extent of fluid-rock

interaction, the composition of fluids (and hence the fluid source), the ambient

temperature and to calculate the fluid flux (volume and rate of fluid infiltration).

42 IntroductionInterpretations of stable isotope data often rely on an assumption that no stable

isotope variation occurs perpendicular to observed isotopic fronts (i.e. one-dimensional

models) and that alteration is the result of a single episode of fluid infiltration. Many of

these assumptions have been challenged by detailed petrological and stable isotope work

in the Cyclades (e.g., Ganor et al., 1996) (where outcrop and grain-scale variations of

isotopic composition have been controlled by selective infiltration of small amounts of

fluid (Bröcker, 1990) and multiple fluid flow episodes have been recognised (Baker et

al., 1989). M2 overprinting of M1 metamorphic assemblages in the Cyclades has been

ascribed to fluid flow (Matthews and Schliestedt, 1984; Bröcker, 1990). The Cyclades

have been an important area in the development of models of fluid-rock interaction and

this will be discussed further in the Chapter 6.

1.13 Application of SHRIMP dating to the Cyclades

One aim of this project was to test the applicability of SHRIMP U-Pb ages to

solving complex geological problems by placing constraints on the timing of

metamorphism and its relationship to magmatic activity during the Alpine orogeny in the

Cyclades. Several avenues of potential research were explored including the development

and application of SHRIMP dating techniques to uranium-bearing accessory minerals

grown during metamorphism, such as titanite. Unfortunately most of the metamorphic

minerals of interest such as titanite and rutile were found to contain very low levels of

uranium (< 10 ppm), making them unsuitable for dating purposes. However, most of the

samples did contain zircon and a few samples contained dateable titanite and monazite.

Zircon, a mineral routinely dated by SHRIMP, was expected to be of little use in

constraining the timing of metamorphism in the area, as most of the Cycladic rocks have

undergone greenschist to amphibolite grade metamorphism at temperatures which did not

exceed 550 ˚C, whereas the limited literature published on the mechanisms by which

zircon is produced, generally describe zircon formation in high-grade partially molten

rocks (migmatites) (Gastil et al., 1967; Marshall, 1969; Gupta and Johannes, 1985) or

granulite-grade rocks (Williams and Claesson, 1987; Kröner and Williams, 1993). Zircon

ages are generally thought to remain unaffected by metamorphism below granulite grade

(Davis et al., 1968). However, some zircons from the Cycladic samples do display

narrow young growth rims around older cores, and the age of these rims is consistent

with the timing of the Alpine orogeny. Hence, not only was the original aim of retrieving

information about the Alpine orogeny achieved but because zircon cores retain their

original age of formation, a property known as “inheritance”, the age of the original

zircon source, or protolith, could also be established.

When zircons record several growth layers of different ages, they are impossible to

date by conventional TIMS which does not provide the spatial resolution required to

distinguish between different-aged growth zones (Williams, 1992). This type of

Chapter 1 43information requires a dating technique that allows within-grain analysis of zircon

structures. SHRIMP was designed for exactly this purpose, and can analyse spots less

than 20 µm in diameter (see Appendix D). As different zircon growth zones can be dated

using SHRIMP, these can be related to different periods of either magmatism or

metamorphism in the zircon’s history thus providing clues to the timing of tectonic events

in the Cyclades. Inherited and detrital zircon ages place maximum constraints on the

depositional age of metasedimentary rocks. Zircon inheritance patterns may also be used

to distinguish the character of the sources from which igneous rocks were derived.

U, Th and Pb/U in the sample zircons were measured according to the procedures

described in Appendix D and the results are listed in Appendix E. Zircons were imaged

both before and after SHRIMP analysis using cathodoluminescence (CL) techniques

(Appendix C) to identify internal zircon structures and to locate exactly the analysis pit

created by the ion beam relative to them. This procedure ensured the recognition of any

mixed ages caused by overlap of the pit across more than one growth zone.

1.13.1 Zircon inheritance patterns

Ages from the twenty-six different samples analyses are combined in order to

define representative age populations for the Basement and Series rocks of the Cyclades.

This information can then be used to characterise tectonically active periods of time in the

area. To illustrate the validity of this approach, the results from one metasedimentary

sample from the Mesozoic Series rocks of Naxos (NX9490) will be described. This

sample shows that many separate age populations can be identified from zircons within a

single sample (Figure 1-23).

44 Introduction

0

1

2

3

4

5

6

0 500 1000 1500 2000 2500 3000 3500

No.

Ana

lyse

s

Age (Ma)

NX9490n = 50

Figure 1-23: Combined plot of age data for 50 analyses of zircons from sample NX9490, ametasedimentary rock from Naxos, utilising a histogram with 50 Ma bin widths overlain by a kernedprobability density curve.

Sample NX9490 zircon ages are presented graphically on combined histograms

overlain by a kerned probability-density curve. This curve allows for unequal analytical

errors, avoids the problem of bin-width encountered in a histogram presentation and

assists in the identification of separate age populations which appear as peaks (Silverman,

1986). Similar graphical representations will be used for age-data from other samples

throughout this thesis. Unless explicitly stated, all ages reported in this thesis are206Pb/238U ages for samples younger than 1000 Ma and 207Pb/206Pb ages for older

samples and the error on all ages are quoted at the ± 1 σ level.

Several hundred samples of pre-Mesozoic basement and Mesozoic series rocks

were collected for a regional study of the geological evolution of the Cyclades. These

units represent a large proportion of the crustal material which crops out in the Cyclades

and should give a good representation of the crustal signature of the area. The U-Pb

compositions of zircons from forty-one samples were measured using the ANU’s

SHRIMP I and II ion microprobes (all results are tabulated in Appendix E). The

youngest zircons yield ages ranging from Miocene to Carboniferous. All of the samples

analysed were post-Devonian in age, thus all pre-Carboniferous ages can be considered as

“inherited” or protolith ages.

Chapter 1 45Applying a test of adequacy (Appendix D, section D10.1) to the data from NX9490

shows that fifty analyses will yield a 95% certainty of finding components in the

population which are over 5.8% abundant. The combined plots in Figure 1-23 shows

that there are scattered ages older than 1000 Ma with peaks at ca . 2450 and 2000 Ma and

sharply defined younger peaks at ca . 800 and 600 Ma. These age populations commonly

occur in zircons from different samples from different Cycladic islands. Younger ages

tend to outnumber older ages presumably due to their better preservation potential in the

geological record, although this also reflects the age of the source material available to

contribute to sedimentation. One analysis of new metamorphic zircon growth is recorded

in this sample at ca . 45 Ma, identified as metamorphic on the basis of its morphology and

low Th/U ratio (see Chapter 6). The next youngest age, considered to be from a zircon

recording the protolith age rather than a metamorphic zircon, occurs at ca . 100 Ma. This

age constrains the time of deposition of the sediment forming sample NX9490 to be

younger than ca . 100 Ma (Cretaceous). Figure 1-24 shows some of the main age

populations encountered in samples from all islands except Folegandros and demonstrates

that although most of the samples were derived from Naxos, many of the age populations

are found in samples from other islands.

++++ ++++ ++ +

100 1000Age (Ma)

Syros

Sifnos

Paros

Naxos

Sikinos

Ios

++ +++++++++ + +++ ++

++

+ + ++++++ +

25 50 200 500

+

Figure 1-24: Schematic representation of the distribution of major zircon age populations for theCyclades on a plot of age to reduce the x-axis scale. Each cross is representative of a population of morethan 5 individual ages.

Different rock-types may yield different zircon age populations, for example, an

orthogneiss will preserve zircons that crystallised during the formation of a pre-existing

magmatic rock, which may in turn preserve inherited cores that will yield information

about the protolith from which the magma was derived. In contrast zircons from a

paragneiss would be expected to be sedimentary in origin, with the ages of these grains

providing information about the source area from which the sediment was derived.

This type of age information is unique and can only be extracted from zircon using

an instrument with a within-grain analysis capability like SHRIMP. This uniqueness is a

46 Introductiondisadvantage in some respects because the data are strictly comparable only to other

SHRIMP age data sets which record multiple age populations from single samples. This

dating study represents the first of its kind conducted in the Cyclades. Information about

the deposition and provenance of sediments in the Cyclades was previously scarce and

reliant on rare palaeontological constraints. There are no similar published age data

available for any part of the southern Alps and only sparse information is available for

other parts of Europe and Africa.

Dating of primary and inherited zircons, combined with age information from

monazite and titanite has proved an exceptional means to address a range of geological

problems in the Cyclades. The geologic history of the area prior to the Alpine orogeny is

crucial in testing plate reconstructions of the area. Due to the complex geological history

and limited outcrop available, it is difficult to develop plate reconstructions prior to the

Alpine orogeny. U-Pb ages on zircon, monazite and titanite can help constrain tectonic

models for the area and facilitates more reliable correlations between the Cyclades and

adjacent terranes.

1.14 Sample Selection

Several hundred samples were collected during the course of this study from both

the Basement and Series rocks of the Cyclades, encompassing a range of lithologies

selected to concentrate U-bearing metamorphic minerals, titanite and rutile. Sampling was

concentrated on these units as, unlike the Upper Unit of the Cyclades, they have

experienced the high-P metamorphism associated with the Alpine orogeny. As mentioned

earlier, most of the titanite and all of the rutile separated from these samples proved

unsuitable for dating purposes because of their exceedingly low uranium contents. Many

of the samples did contain zircon, and samples from the high grade core of Naxos also

contained monazite, both of which proved readily dateable. Thus, although samples were

collected from most islands of the Cyclades and from diverse lithologies and metamorphic

grades, the samples described in this thesis represent only those which have yielded

dateable minerals (41 different rock units in total).

Inherited zircon ages from twenty-six samples of both Basement and Series rocks

from seven islands, Naxos, Paros, Ios, Syros, Sifnos, Sikinos and Folegandros, have

been obtained. Their relevance to correlations between the Cyclades and other areas are

discussed in the next chapter. To characterise the time of formation of the Cycladic

basement, zircons have been dated from seventeen samples from four islands from

material distinguished as Cycladic basement by previous workers. The results are

described in Chapter 3 and include six samples from Ios, two samples from Paros, one

sample from Sikinos and eight samples from Naxos. Similarly the Cycladic Series rocks

have been characterised by dating of zircon from nineteen samples from six islands

Chapter 1 47including nine samples from Naxos, four samples from Syros, three samples from Ios,

one samples from Folegandros, one sample from Sikinos and one samples from Sifnos.

The zircon protolith ages derived for the Series rocks are described in Chapters 3 and 4.

The ages of metamorphic monazite, titanite and zircon from samples from Naxos and

Sifnos are described in Chapter 6. The timing of Miocene magmatism is constrained by

dating of zircon, titanite and monazite from eight samples from Naxos and one sample

from Tinos, with results described in Chapter 7. Chapter 8 combines all of the age data

gathered to yield a coherent picture of the geological evolution of the Cyclades from the

Archaean to the present.

Chapter 2 49

2. PRE-CARBONIFEROUS EVOLUTION OF THE CYCLADES

2.1 Introduction

Dating detrital minerals in sediments is a common method used to determine

sedimentary provenance (Hurley et al., 1961), to define periods of crustal growth

(Tatsumoto and Patterson, 1964) and to trace the origin of allochthonous terranes (Krogh

and Keppie, 1990). Similar information about crustal evolution can be gained by dating

inherited zircon grains in granites (Pidgeon and Aftalion, 1978) and orthogneisses

(Compston and Kröner, 1988) and information about the source regions from which

granites were derived (Miller et al., 1992) (see Appendix A). Detrital and inherited

zircons from both the Basement and Series rocks of the Cyclades have been dated in this

study to provide information about the early history of Cycladic crust.

This is the first study to constrain the early geological evolution of the Cyclades by

U-Pb dating of inherited and detrital zircons to determine provenance. The previous

identification of pre-Mesozoic basement in the Aegean (van der Maar et al., 1981; Henjes-

Kunst and Kreuzer, 1982; Andriessen et al., 1987) provides a starting point for the early

geological history of the Cyclades, but age information from this basement is generally

confined to the post-Carboniferous and is discussed in Chapter 3. In pre-Mesozoic times,

the continental blocks of the Hellenides including the Attic-Cycladic Massif (containing

the Cyclades) are thought to have formed part of the northern margin of Africa (Smith,

1971; Dewey et al., 1973; Aubouin, 1976; Sengör and Yilmaz, 1981; Robertson and

Dixon, 1984; Sengör, 1984; Sengör et al., 1988) (Figure 2-1).

50 Pre-Carboniferous

EASTGONDWANA

PACIFIC OCEAN

0o

30o

60o

BALTICA

LAURENTIA

Cadomian arc

~ 545 Ma

South Pole

WESTGONDWANA

Cyclades

Figure 2-1: Geologically constrained reconstruction for the Precambrian-Cambrian boundary showingthe distribution of major continents (from Dalziel, 1997). The position of the Cyclades on the northernmargin of the African continent is highlighted.

Plate reconstructions for the pre-Mesozoic are particularly difficult (Dercourt et al.,

1986), given that the outcrop patterns have been disrupted by the Variscan and Alpine

events, and the palaeontological and palaeomagnetic controls in many areas such as the

Cyclades are absent (Dürr et al., 1978; Morris and Tarling, 1996). However, the general

location of the major continental blocks is fairly well-constrained and the Precambrian

geological evolution of the Cyclades is expected to be broadly similar to that of northern

Africa (Figure 2-1).

This chapter reports pre-Carboniferous U-Pb inherited and detrital zircon ages from

orthogneisses and metasediments collected from seven of the Cycladic islands. Several

distinct age peaks have been identified which may assist in comparisons with other areas

Chapter 2 51of the Hellenides and Taurides and provide clues to links with other areas within the

Alpine orogenic belt.

2.2 Geologic Setting

Correlations between the Pelagonian zone, the Menderes Massif and the Cyclades

have been suggested due to the geological similarities between these terranes and their

close spatial relationship (Dürr et al., 1978) (Figure 2-2). As discussed in subsequent

chapters, both the Basement rocks and the Series cover sediments in the Cyclades are

post-Carboniferous in age, but both lithological groups contain inherited zircons that limit

the pre-Carboniferous history. Between them, the two groups comprise a large

proportion of the crustal material in the Cyclades, and they are considered to have

inherited a zircon component that is representative of the pre-Carboniferous crustal

signature of the area.

TURKEY

GREECE

MenderesMassif

Pelagonian Zone

Cyclades

External Hellenides

Figure 2-2: Structure of the Eastern Mediterranean showing the extension of the Attic-Cycladic Massifinto Turkey (the Menderes Massif) and Greece (the Pelagonian zone).


Syros

Naxos

Folegandros

Sikinos

Ios

Sifnos

Paros

0 20km

N

EW

S

CYCLADES

Upper UnitMiocene IntrusivesMesozoic SeriesPre-Alpine Basement

Sample Location

Figure 2-3: Location of samples with pre-Carboniferous zircons. See Appendix B for more sampledetails and Appendix E for sample numbers and details of analytical results.

2.3 Previous Geochronology

Sparse information about the pre-Carboniferous history of the Cyclades has been

obtained from dating of pre-Mesozoic basement yielding U-Pb, Rb-Sr and K-Ar ages in

the range 1000 - 300 Ma (van der Maar et al., 1981; Henjes-Kunst and Kreuzer, 1982;

Andriessen et al., 1987). Conventional U-Pb dating of zircons from Ios basement

yielded a linear array with a lower intercept at 305 - 300 Ma and an upper intercept ~ 1000

Ma on concordia (Henjes-Kunst and Kreuzer, 1982). The lower intercept age was

interpreted as reflecting strong episodic Pb loss associated with recrystallisation of the

orthogneisses during high grade Variscan metamorphism (see Chapter 3), while the upper

Chapter 2 53intercept age was taken as a minimum estimate of the age of the inherited component in

the zircons. A zircon crystallisation age of 372 Ma (+28/-24) was obtained from the

migmatite core of Naxos by conventional U-Pb dating (Andriessen et al., 1987). This

age was considered to be within error of a ~ 500 Ma Rb-Sr isochron age from Ios

orthogneisses (Henjes-Kunst and Kreuzer, 1982) and was used to suggest the granite

protoliths were intruded between the Cambrian and Devonian (although the results from

this study presented in Chapter 3 indicate a Carboniferous emplacement age for the

orthogneiss precursors on Ios, Paros and Naxos).

2.4 U-Pb Analytical Results

Twenty-six samples yielded pre-Carboniferous ages (Table 2-1) either from detrital

zircons in metasediments or from inherited cores in magmatic zircons from orthogneisses.

The samples included eighteen metasediments and eight orthogneisses from seven

different islands; Folegandros, Ios, Naxos, Paros, Sifnos, Sikinos and Syros (Figure 2-

3).

Table 2-1: Samples that have yielded pre-Carboniferous ages.

Island Sample Rock-type Unit Ages(> 355 Ma)

Folegandros FL9602 Pelite Series 16Ios IO9606 Garnet-mica schist Basement 11

IO9607 Leucogneiss Basement 8IO9609 Garnet-mica schist Basement 7IO9615 Garnet-glaucophane schist Series 189640 Orthogneiss Basement 289639 Glaucophane schist Series 190346 Quartz-phengite schist Series 3IO9403 Orthogneiss Basement 8IO9404 Orthogneiss Basement 3

Naxos NX9314 Orthogneiss Basement 3NX94106 Pelite Series 3NX94112 Calc-silicate Series 1NX94121 Calc-silicate Series 8NX9461 Calc-silicate Series 2NX9463 Calc-silicate Series 10NX9464 Calc-silicate Series 32NX9485 Orthogneiss Basement 1NX9490 Pelite Series 47NX9638 Migmatite Basement? 7

Paros PA9601 Orthogneiss Basement 1PA9606 Orthogneiss Basement 2

Sifnos SIF9345 Calc-silicate Series 13Sikinos SK9601 Orthogneiss Basement 4Syros 89646 Quartzite Series 7

Representative images of the different zircon morphologies encountered in this

study are shown in Figure 2-4. A large range in zircon types can be recognised

petrographically, from dark-coloured, abraded, detrital zircons to clear, elongate, euhedral

54 Pre-Carboniferouszircons which do not appear to have undergone a sedimentary cycle. Old inherited

sedimentary zircon grains can be distinguished petrographically by their pitted surfaces,

indicating abrasion during transportation, or by their occurrence as xenocrystic cores

(identified using CL images) overgrown by younger oscillatory-zoned rims. Old detrital

grains and inherited cores are the most common type of zircon in the metasedimentary

samples and contributed to the older age populations identified. The orthogneisses

contain more zircons of magmatic appearance, having elongate grains that display well-

developed crystal faces with sharp terminations and that exhibit oscillatory growth zoning

in CL. These zircons tend to yield relatively young (< 600 Ma) ages. Zircons from the

metasediments contain a higher proportion of inherited cores than zircons from the

orthogneisses.

PA960689646a.

200 µm

b. NX9490

200 µm

d. IO9609

200 µm

c. PA9601

200 µm

Figure 2-4: Transmitted light photomicrographs of zircons from samples: a) 89646, a quartzite fromSyros with abraded detrital grains; b) NX9490, a low M2 grade pelite from Naxos with abraded detriatlgrains; c) PA9601, an orthogneiss from Paros with elongate magmatic-appearing grains and; d) IO9609, agarnet mica schist from Ios showing a range of zircon morphologies.

The ages from all twenty-six samples have been combined and a summary of major

zircon age components older than Carboniferous is given in Table 2-2, whereas individual

analytical results are listed in Appendix E. The age populations are graphically displayed

in Figure 2-5. Scattered peaks are shown in the Palaeoproterozoic and Archaean,

reflecting the relatively small number of analyses (n = 26 and 12 respectively). The two

oldest ages measured at ca . 3170 and 3190 Ma were from the metasedimentary sample

NX9490 from Naxos and from a metasedimentary sample from Folegandros. Several

Chapter 2 55peaks can be identified at ca . 2900-2850, 2500-2450, 2050-2000, 1900-1800 and 1700-

1650 Ma. Of these Palaeoproterozoic/Archaean ages, sixteen of the thirty-eight analyses

were from obvious xenocrystic cores while the other analyses generally came from

irregularly zoned grains.

Table 2-2: Zircon Age Components

Era Age Population (Ma)

Palaeozoic 450-400

625-525

Neoproterozoic 675-625

900-800

1000-950

1700-1650

Palaeoproterozoic 1900-1800

2050-2000

2500-2450

Archaean 2900-2850

0

5

10

15

20

25

30

0 500 1000 1500 2000 2500 3000 3500

P NP MP PP A

No.

Ana

lyse

s (n

)

Age (Ma)

Figure 2-5: Pre-Carboniferous age distribution for the Cyclades from 202 zircon analyses. The graphis comprised of a histogram of age data using 50 Ma bin widths overlain by a kerned probability densitycurve to illustrate the major age peaks. P - Palaeozoic, NP - Neoproterozoic, MP - Mesoproterozoic, PP- Palaeoproterozoic, A - Archaean (timescale from the AGSO Phanerozoic Timescale of Young andLaurie, 1996).

56 Pre-CarboniferousThere is a distinct lack of Mesoproterozoic ages in the zircon populations. Only ten

zircons yielded ages in the range 1600-1000 Ma and these were from metasedimentary

units on Naxos, Ios and Folegandros. A significant age gap is apparent from ca . 1600-

1400 Ma. There is a complicated distribution of ages between 1100-700 Ma with two

peaks distinguishable at ca . 1000-950 and 900-800 Ma. Distinct peaks occur at ca . 650

and 550 Ma while there is one pre-Carboniferous Palaeozoic peak in the range ca . 450-

400 Ma. The peaks are all defined by more than twenty individual analyses, with ages

less than 650 Ma comprising over half (n = 114) of the total number of analyses (n =

202).

2.5 Discussion

The dominant inherited zircon age populations visible in Figure 2-5 and listed in

Table 2-2 can be used to characterize the crustal precursors from which Basement and

Series rocks of the Cyclades were ultimately derived. Unfortunately, little work has been

published on inherited zircon age populations for the Alpine orogens, making correlations

between areas difficult. Some ion-probe analyses of pre-Mesozoic detrital, metamorphic

and igneous zircons from a range of rock-types (Gebauer et al., 1989; Gebauer, 1993)

have been used to constrain the pre-Mesozoic evolution of Central European continental

crust. This crust, like that of the Cyclades, is thought to have been derived from the

northern margin of Gondwana because of the predominance of Pan-African ages found.

The Pan-African has been defined very broadly as any crustal material aged between 950-

450 Ma (Kröner et al., 1987), but a more recent definition will be used in this study that

restricts ages to the range 730-550 Ma (Black and Liegeois, 1993).

Gebauer (1993) distinguished four Precambrian “mega-cycles” of crustal

development in central Europe at ca . 2700-2500, 2200-1900, 1200-900 and 800-550 Ma

and interprets the abundance of Pan-African ages as evidence that the European Variscides

(central Europe) was derived from the supercontinent Gondwana rather than Laurasia. In

contrast, the Laurasian crust of the Caledonides yields Mesoproterozoic ages (Williams

and Claesson, 1987; Stephens et al., 1993), which are lacking in the age populations of

central Europe and the Cyclades. The Mesoproterozoic age gap in the Cyclades is

consistent with its derivation from North Africa and reflects a lack of material of this age

in the area (Cahen et al., 1984; Goodwin, 1995) (Figure 2-6).

This absence of Mesoproterozoic ages is also distinctive of West African crust and

has been used in a similar way in the Appalachians to distinguish between accreted

terranes derived from the South American and the West African portions of West

Gondwana (Nance and Murphy, 1994) (Figure 2-1).

Chapter 2 57

Figure 2-6: Distribution of Precambrian crust of different ages across the continents joined in a pre-driftPangean reconstruction (adapted from Goodwin, 1991). Note the lack of Mesoproterozoic ages in NorthAfrica.

Gebauer (1993) distinguished the period between ca . 1700-1200 Ma as one of

tectonic quiescence in Central Europe and suggested that this is typical for the “non-

Australian” part of Gondwana, presumably West Gondwana. A “Gondwana component”

containing Proterozoic age peaks at ca . 1700-1400 and 1150-1000 Ma is shared by

Palaeozoic sediments from New Zealand, Eastern Australia and Antarctica (Ireland, 1992;

Wysoczanski et al., 1997). The occurrence of Mesoproterozoic zircon ages, which

presumably record important tectonomagmatic events in segments of East Gondwana and

also in the Amazonian craton (Bernasconi, 1987; Teixera et al., 1989) suggests an

important distinction can be made between zircon populations of these areas and those of

North and West Africa.

Neoproterozoic ages in the range 900-700 Ma are considered to be characteristic of

North African and Arabian crust (Kröner and Sengör, 1990), where these ages are

derived from primitive magmatic rocks produced during the development of an active

continental margin and associated island arcs (Stern, 1981; Reischmann et al., 1992;

58 Pre-CarboniferousStern and Kröner, 1993). The identification of a significant Neoproterozoic age

population in this study supports a close relationship between the Cyclades and North

Africa, where rocks in this age range are common (Cahen et al., 1984) (Figure 2-6), and

is consistent with proposed tectonic reconstructions of the area based on palaeontological

and lithological constraints (e.g. Robertson and Dixon, 1984). The latter would also

place the Menderes Massif of western Turkey in close proximity to the Cyclades as part of

the northern margin of Gondwana. The model is supported by the similarities in

Precambrian age patterns between the Cyclades and those reported for the Menderes

Massif from conventional U-Pb zircon and single zircon Pb evaporation, which span ca .

2555-1740 and 1000-700 Ma (Reischmann, 1991). However, Kröner and Sengör,

(1990) disputed a connection between western Turkey and north Africa on the basis of

zircon Pb evaporation ages which exhibit a gap in the 900-700 Ma age range for samples

from the Menderes-Taurus block of southern Turkey. Kröner and Sengör, (1990),

suggested that this precludes derivation of the Menderes Massif from Africa and proposed

that it was derived instead from the East Sayan block of the Angara craton of Siberia.

While the mega-cycles recognised in Central Europe are broadly similar to the age

components found in the Cyclades, they differ in one important respect. The Central

European samples do not appear to contain a dominant bimodal age population in the

restricted ca . 650-550 Ma range. Many workers would consider the latter to represent the

“true” Pan-African orogenic cycle, with such ages widespread through all segments of

Gondwana (Figure 2-6 and Figure 2-1) although being given different names e.g.

Brasiliano (South America), Delamerian (Australia), and Beardmore (Antarctica). The

zircon age patterns for the Cyclades likewise distinguish two separate Pan-African events,

one at ca . 650 Ma and one at ca . 550 Ma. The two age populations are commonly found

in East Africa and are thought to represent two distinct orogenies: the East African

orogeny (800-650 Ma, Stern, 1994) and the Kuunga orogeny (550-530 Ma, Meert et al.,

1995). One explanation for the two ages is that the ca . 650 Ma peak represents the

suturing of East and West Gondwana while the 550 Ma peak reflects minor adjustments

within the assembled supercontinent (Unrug, 1992; Stern, 1994) or the possible

extensional collapse of the East Africa Orogen (Windley et al., 1994). Another

explanation based on palaeomagnetic data is that Gondwana was finally assembled during

the Kuunga orogeny and that the East African Orogeny predated this event (McWilliams,

1981; Li and Powell, 1993; Powell et al., 1993). The Kuunga orogeny was a time of

widespread granulite formation throughout East Gondwana (Meert and van der Voo,

1997), and the 550 Ma ages may correspond to a second collisional event between

Australo-Antarctica and the rest of Gondwana which had already been amalgamated by

650 Ma (Meert et al., 1995).

Chapter 2 59Pan-African ages are also common in both detrital and magmatic zircons from the

Menderes Massif of western Turkey (Kröner and Sengör, 1990; Kampunzu and Lubala,

1991; Loos and Reischmann, 1995; Hetzel and Reischmann, 1996), which would

support the idea of an affinity between the rocks of western Turkey and the Cyclades. No

ages in the 900-700 Ma age range have been reported for western Turkey, although these

ages define a distinct peak in the Cyclades. The significance of this difference in age

populations is unclear, and may be resolved by a more detailed study of zircon inheritance

patterns from the Menderes Massif. The increased frequency of Pan-African zircons in

the eastern Mediterranean region, as found in this study, may assist with the delineation

of a belt defined by Pan-African tectonism at the margins of the Central European craton.

However, there are few ages in the Menderes Massif which match the strong 450-400 Ma

age group found in the Cyclades and the Menderes Massif is thought to have remained

relatively stable from the early Ordovician to at least the late Triassic (Sengör, 1984) with

little tectonic activity noted for the Silurian in Turkey (Brinkmann, 1976). Late

Ordovician orogenic activity is widely documented in the Central Alps with granite

intrusion at ca . 460-440 Ma (Köppel et al., 1980), possibly related to collision of terranes

separated from Gondwana after the Pan-African orogeny (Gebauer, 1993). The absence

of 450-400 Ma ages in western Turkey suggests that while the protoliths from which the

Cyclades were derived experienced similar Precambrian histories to the rocks of the

Menderes Massif, they have experienced distinctly different Palaeozoic histories. This

should be an important consideration in plate reconstructions of the eastern Mediterranean

and is not adequately addressed by current models (e.g., Robertson and Dixon, (1984);

Kröner and Sengör, (1990).


2.6 Synthesis

Zircons from Palaeozoic sediments and orthogneisses from the Cyclades record

complex growth histories that provide information about the source rocks from which

these sediments and intrusives were derived. Zircon age patterns for the Cyclades and

western Turkey show similarities for the Precambrian but distinct differences in the

Palaeozoic. The lack of published pre-Mesozoic zircon age data for other areas precludes

detailed comparisons being made with the Cyclades. Nevertheless it can be concluded

that the Cycladic age pattern is similar to that reported for the central Alps, and that the

preponderance of Pan-African ages suggests affinities with Gondwana rather than

Laurasia, consistent with interpretations that the Cyclades were derived from the northern

margin of Gondwana (presently North Africa). The distinctive Mesoproterozoic age gap

found in zircons from Cycladic rocks is also found in the West African craton and is a

feature which should prove characteristic of North African crust. This Mesoproterozoic

age gap makes it possible to distinguish crust derived from North and West Africa from

that derived from other parts of Gondwana.

Chapter 2 61

Chapter 3 61

3. PERMO - CARBONIFEROUS GEOLOGICAL EVOLUTION OF

THE CYCLADES

3.1 Introduction

The age and character of the Basement rocks of the Cyclades and the extent to

which they have been affected by the Variscan orogeny are investigated in this chapter.

The stable forelands of Eurasia, Africa and Arabia are thought to be underlain by a

continous continental basement of pre-Triassic age which is exposed in many areas of the

Alpine orogenic system (Dewey et al., 1973). Although the age of the earliest formed

rocks in the Cyclades is not well known, Cycladic basement has been likened to similar

orthogneiss associations found in the Pelagonian zone of mainland Greece, the external

Hellenides exposed on Crete and the Menderes Massif of Turkey (Dürr et al., 1978; Blake

et al., 1981; Okrusch et al., 1984). In the Cyclades the orthogneisses and their hosts

preserve evidence of a pre-Alpine metamorphic (M0) and structural (D0) history (Section

3.2) (van der Maar et al., 1981; Henjes-Kunst and Kreuzer, 1982; van der Maar and

Jansen, 1983; Buick, 1991; Franz et al., 1993; Grütter, 1993). The timing of the pre-

Alpine tectonic history is also poorly known, with best estimates suggesting it is Variscan

in age (Henjes-Kunst and Kreuzer, 1982) (Section 3.3). The Variscan orogeny can be

broadly correlated to the Early Palaeozoic (Weber, 1984) (see Chapter 1), with most

magmatic activity restricted to the Permo-Carboniferous (Bonin et al., 1993).

In Permo-Triassic time the two supercontinents Gondwana and Laurasia were

separated by a gulf of water known as the Tethys (Suess, 1875). Suturing of these

continents and closure of the Palaeozoic (Palaeo-) Tethys, resulted in formation of the

Variscan fold belt in the core of the supercontinent (Figure 3-1) (Ziegler, 1993). The fold

belt eventually collapsed as the result of widespread extensional tectonism during the late

Carboniferous (Rey et al., 1991; Reinhardt and Kleeman, 1994; Costa and Rey, 1995).

Subsequently, continued movements between the African and Eurasian plates culminated

in a series of collisional events to form the Alpine orogenic belt (Dewey et al., 1973)

discussed in detail in Chapter 6. The Alpine orogeny has obscured the pre-existing

boundaries of the Variscan belt. Though the southern boundary of the Variscan fold belt

is poorly defined, voluminous Permo-Carboniferous granitoids associated with late stages

of the orogeny have been identified by U-Pb geochronology and used to infer the extent

of Variscan outcrop. These granitoids have intruded the central Alps (Schaltegger and

Corfu, 1992), Dinarides, external Hellenides (Pe-Piper, 1982), internal Hellenides

(Yarwood and Aftalion, 1976; Mountrakis, 1984), Taurides, Sardinia (Neubauer and

Raumer, 1993) and north-east Africa (Cahen et al., 1984). The magmatism has been

variously described as the result of post-collisional orogenic collapse and delamination

62 Permo-Carboniferous(Dewey, 1988; Menard and Molbar, 1988) or Andean-type subduction of oceanic crust

(Mercolli and Oberhänsli, 1988; Finger and Steyrer, 1990; Schaltegger and Corfu, 1992).

An overview of Late Variscan magmatism preserved in the Alpine chain has been

presented by Bonin et al. (1993).

Equator

AustraliaAntarctica

India

Siberia

Africa

EARLY PERMIAN ~280 Ma

Panthalassa

LAURASIA

Variscan

GONDWANASouth

America

Cyclades

Figure 3-1: Early Permian plate reconstruction showing major continental blocks assembled to formthe supercontinent Pangea with the extent of the Variscan fold belt delineated by diagonal stripes (adaptedfrom Ricou, 1994). The Cyclades are labelled forming part of the complex margin of north Africa.

3.2 Geological Background

The Basement rocks of the Cyclades are considered to be pre-Mesozoic in age on

the basis of their structural position below the Series sediments, by the fact that they

preserve structures which are not found in the Series rocks and from radiometric dating

(van der Maar et al., 1981; Henjes-Kunst and Kreuzer, 1982; Andriessen et al., 1987).

The rocks consist mainly of deformed and metamorphosed intrusives and a complicated

sequence of schists, gneisses and amphibolites. Van der Maar et al. (1981) likened the

meta-intrusives of Sikinos to the Ios basement, while van der Maar and Jansen (1983)

postulated the existence of such a basement on Naxos, Mykonos, Paros, Serifos and

Syros. It has been suggested that the Basement has undergone a Variscan amphibolite-

facies metamorphism (M0) with reports of the existence of relict mineral assemblages in

Basement garnet-mica schists which are not characteristic of the Alpine (M1) blueschist

facies metamorphism (Henjes-Kunst and Kreuzer, 1982). These assemblages testify to

the existence of an older amphibolite facies metamophism (M0) with metabasites

containing brown-green hornblende + diopside + plagioclase and metapelites containing

old almandine garnet and pseudomorphs after staurolite. Evidence of M0 is recorded in

Sikinos basement schists by the occurrence of pseudomorphs after staurolite and brown

magnesio-hornblende, thought to indicate P-T conditions of amphibolite facies (~ 5 kbars,

Chapter 3 63570-650 °C (Franz et al., 1993). Grütter (1993) suggested that the basement on Ios

contains an undescribed planar and linear fabric (D0) that is probably Variscan in age.


There are two primary radiometric studies of the orthogneisses comprising the

Cyclades Basement rocks. Both suggest that the protoliths to the orthogneisses on these

were emplaced at some time during the Palaeozoic, although they are limited by the dating

techniques employed. Rb-Sr and U-Pb results from the Ios basement reveal Variscan

ages thought to chronicle the M0 metamorphism of the orthogneisses (Henjes-Kunst and

Kreuzer, 1982). A Rb-Sr whole rock isochron of approximately 500 Ma was obtained

from relict intermediate intrusions (462 ± 48 Ma from eight samples, 520 ± 55 Ma with

three samples removed). Since these intrusions have largely escaped polyphase

deformation, preserving primary igneous textures and mineralogies, the 520 Ma Rb-Sr

age was interpreted as an original igneous emplacement age. Conventional U-Pb dating

of zircons from two samples of similar lithology produced a cluster of ages near 320 Ma

on Concordia. Seven zircon fractions from one sample yielded a linear array with lower

intercept at 305-300 Ma (± 5) and upper intercept ~ 1000 Ma. The lower intercept age

was interpreted as reflecting strong episodic Pb loss associated with recrystallisation of

the orthogneisses during high grade Variscan metamorphism, whereas the upper intercept

was taken as a minimum age estimate of the inherited component in the zircons. The

zircon ages are in agreement with Rb-Sr muscovite-whole-rock(WR) isochron ages of

295 ± 4, 294 ± 4 and 288 ± 8 Ma for three relic intrusives, which were interpreted to

reflect cooling after the M0 Variscan metamorphism (Henjes-Kunst and Kreuzer, 1982).

Andriessen et al. (1987) concurred with the interpretation of Henjes-Kunst et al.

(1982) with regard to the history of the pre-Mesozoic basement in their own U-Pb, Rb-Sr

and K-Ar investigation of the Cyclades. They obtained a zircon crystallisation age of 372

Ma (+28/-24) from the migmatite core of Naxos by conventional U-Pb dating after coarse

zircon fractions were removed to minimise the effects of inherited radiogenic Pb. This

age was considered to be within error of the ~ 500 Ma Rb-Sr isochron age from Ios

orthogneisses (Henjes-Kunst and Kreuzer, 1982). Schuiling (1973) reported a Rb-Sr

whole rock age of 355 Ma from the Naxos migmatite. Andriessen et al. (1987) also

reported a K-Ar hornblende age of 268 ± 27 Ma from the Ios basement which they

interpreted as a post-M0 cooling age, and on Sikinos, Rb-Sr whole-rock measurements

from five samples of a metadiorite intrusion produced a crude linear array yielding an

isochron of 275 ± 87 Ma. These previous investigations all support the existence of pre-

Mesozoic basement in the Aegean, intruded by Palaeozoic granitoids.

64 Permo-Carboniferous3.4 SHRIMP U-Pb Results

Seventeen zircon-bearing gneiss and migmatite samples from areas identified as

basement units in the Cyclades were analysed in this study. Four islands were

investigated: Naxos (8 samples), Ios (6 samples), Sikinos (1 sample) and Paros (2

samples). Two of the six samples from Ios were from metasediment (garnet-mica schist)

hosting the orthogneisses of the Ios basement. There are no previous age constraints on

these schists which represent the country rock that was originally intruded by the

orthogneiss protoliths.

A summary of the U-Pb ages are listed in Table 3-1 (p. 84) with full details of

analytical method for zircon analysis described in Appendix D and all analytical results

provided in Appendix E. As elsewhere in this thesis, the data are also presented in three

graphical formats for each sample: histogram presentations of all ages are overlain by

kerned probability-density plots (i.e. without subtraction of common Pb) and an inset of

the measured 207Pb/206Pb and 238U/206Pb ratios on standard Tera-Wasserburg

Concordia diagrams for each analysis. The figures also show data, where present for

inherited pre-Carboniferous zircon (discussed in Chapter 2) and young zircon reflecting

the Alpine metamorphic overprint (discussed in Chapters 5 and 6).

3.4.1 Ios

The Basement of Ios is exposed over a

large proportion of the island (Figure 3-2) and

consists of an orthogneiss core, mantled by an

envelope of garnet-mica schist, augengneiss and

intrusive intermediate-mafic rocks (Henjes-

Kunst and Kreuzer, 1982; van der Maar and

Jansen, 1983; Andriessen et al., 1987). The

geochemistry of the orthogneisses on Ios

(Henjes-Kunst and Kreuzer, 1982) suggests

they are S-type granites (Chappell and White,

1974), probably derived from a peraluminous

sedimentary protolith. Unlike the Series rocks

of the Cyclades, the Basement contains no

marble units.

Six samples from Ios were analysed to characterise the age of the granites and

sediments that comprise the Basement. These samples include three orthogneisses, one

leucogneiss and two of the enclosing metasedimentary units (garnet-mica schists) which

the orthogneiss protolith is thought to intrude. Locations of the samples are given in

0 2 4 kmN

OrthogneissGarnet-Mica Schist

MarbleSchist

IO9403IO9404 89640

IO9607

IO9609

IO9606

Figure 3-2: Sample location map for Ios(adapted from van der Maar and Jansen, 1981).

Chapter 3 65Figure 3-2. The samples will be described in the following sections based on their

lithologies.

3.4.1.1 Ios Orthogneisses

Investigation of orthogneiss chemistry by Henjes-Kunst and Kreuzer (1982)

suggests that all the intrusives are S-type granites as defined by Chappell and White

(1974). The determination of emplacement ages for S-type granites is commonly difficult

due to the dominance of inherited zircon grains (as illustrated in the Keay et al. paper

listed in Appendix A). However, there is usually some evidence of new zircon growth,

even in S-type granites where crystallised melt appears subordinate to restite material.

The Ios orthogneisses actually contain very little restite, according to the definition of

Chappell et al. (1987), and it seems likely that the youngest zircon populations grew

within the granite magma. Zircons from all three samples were generally euhedral and

elongate, oscillatory-zoned grains of magmatic appearance up to 200 µm in length with

rare irregular-shaped inherited cores, generally identifiable by a truncation in the zircon

growth zoning. Age data for the Ios orthogneiss samples IO9403, IO9404 and 89640 are

given in Figure 3-3, Figure 3-4 and Figure 3-5 respectively. Sample 89640 is from the

RSES, ANU rock collection (collected by Dr. S. Baldwin).

IO9403n = 26

0.20

0.16

0.12

0.08

0.04

0 5 10 15 20 25 30 35 40

200 Ma3004005001000

2000

238 U / 206 Pb

207

Pb /

206

Pb

0

5

10

15

20

200 400 600300 500 700

No.

of

Ana

lyse

s

Age (Ma)

Figure 3-3: Combined histogram with 20 Ma bin widths and kerned probability density curve forzircons from sample IO9403. Inset is a Tera-Wasserburg Concordia diagram showing all analyses. Notethat one analysis at 1090 Ma (207Pb/206Pb age)is not included on the histogram or age probability curve,and that the common Pb correction method projects the radiogenic points exactly on Concordia (207-method).

66 Permo-CarboniferousAll samples show distinct age peaks in the range 330-300 Ma, with minor traces of

inheritance dating back to 1090 Ma (207Pb/206Pb age) No new zircon growths younger

than ca . 292 Ma were discovered, suggesting that the ca . 330-300 Ma ages reflect the time

of emplacement and cooling of the original granite protolith of the orthogneisses.

IO9404n = 8

0.04

0.05

0.06

0.07

0.08

0.09

0.10

0.11

0.12

0 5 10 15 20 25 30

1000

500300 Ma400

238 U / 206 Pb

207

Pb /

206

Pb

0

1

2

3

4

5

300 400 500 600 700 800Age (Ma)

No.

of

Ana

lyse

s

Figure 3-4: Combined histogram with 25 Ma bin widths and kerned probability density curve forzircons from sample IO9404. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

The orthogneisses contain the M1 assemblage garnet-phengite-albite-quartz, which

has replaced most of the original igneous feldspar, although some evidence of igneous

texture is preserved despite the development of a strong foliation. The small number of

inherited cores (16 of the 47 grains analysed) within the 330-300 Ma zircon population

and the generally magmatic appearance of grains (Figure 3-9, Figure 3-13) suggests that

the zircon populations are not entirely inherited. S-type granitoids, by definition, are

derived from sedimentary protoliths and any zircons from such a source would have

undergone a sedimentary cycle which generally causes abrasion and rounding of zircon

grains. A range of inherited ages, similar to those found in the garnet mica schists of Ios

(Section 3.4.1.2), would be expected from the zircons if there was little growth during

melting. Instead, the zircons from the 330-300 Ma orthogneisses are uniformly elongate,

prismatic grains with sharp terminations typical of most magmatic rocks. Preservation of

“pre-Alpine” structures in the basement rocks of Ios (Henjes-Kunst and Kreuzer, 1982;

Grütter, 1993) also support the interpretation that these granites were emplaced at ca . 330-

300 Ma and are not merely deformed Miocene intrusives (McGrath et al., 1995; Pe-Piper

Chapter 3 67et al., 1997). The possibility that these zircons might have undergone recrystallisation

under amphibolite-facies metamorphic conditions as suggested by Henjes-Kunst and

Kreuzer (1982) is discounted, as the zircons show no morphological or textural evidence

of recrystallisation (cf. Pidgeon, 1992).

89640n = 13

0.04

0.08

0.12

0.16

0.20

0 5 10 15 20 25 30 35 40

1000

500 300 200 Ma400

238 U / 206 Pb

207

Pb /

206

Pb

0

1

2

3

4

5

6

250 300 350 400 450 500

Age (Ma)

No.

of

Ana

lyse

s

Figure 3-5: Combined histogram with 10 Ma bin widths and kerned probability density curve forzircons from sample 89640. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

Interestingly, sample IO9404 is an undeformed granite boudin contained within the

IO9403 orthogneiss (samples taken less than 2 metres apart). The strong Alpine

deformation apparent in samples 89640 and IO9403 seems to have had little effect on the

zircon populations, although zircons from the deformed granite IO9403 are generally

smaller in size than those from the undeformed boudin IO9404, suggesting zircon may

have dissolved during Alpine deformation rather than precipitating.

3.4.1.2 Ios Leucogneiss

IO9607 appears to be a metamorphosed aplitic dyke, with an M1 mineral

assemblage of quartz-albite-phengite, which cuts the Ios paragneisses. This dyke yielded

a young Rb-Sr WR-phengite tie-line of 13.2 ± 0.4 Ma, interpreted as an indication of a

Miocene thermal influence (Henjes-Kunst and Kreuzer, 1982; Baldwin and Lister, 1998).

The zircons in this sample are generally small, less than 150 µm in diameter, and contain

68 Permo-Carboniferousobvious inherited zircon cores. These cores commonly comprise a large proportion of the

zircon grain and are overgrown by oscillatory zoned rims (Figure 3-9). Thirteen analyses

of eleven zircon grains yielded a range in ages from ca . 2440 to 295 Ma (Figure 3-6).

Applying Dodson’s test of adequacy (Appendix D10.1) shows that from thirteen

analyses, a 95% confidence level requires measuring an age population which is 21%

abundant in the sample. Three ages form a population at ca . 321 Ma while there are also

two ages at ca . 297 Ma (Table 3-1, Figure 3-6), implying that the aplite might be related

to the emplacement of the Ios orthogneisses between 330-300 Ma. As there are no ages

younger than ca . 295 Ma, the small population at ca . 297 Ma is taken to represent the time

of emplacement of this sample, although more analyses are required to confirm this

interpretation.

IO9607n = 13

0.04

0.06

0.08

0.10

0.12

0.14

0.16

0.18

0.20

0 5 10 15 20 25 30

2000

1000

500 400 300 Ma

238 U / 206 Pb

207

Pb /

206

Pb

0

1

2

3

0 500 1000 1500 2000 2500Age (Ma)

No.

of

Ana

lyse

s


3.4.1.3 Ios Garnet Mica Schists

Two samples (IO9606, IO9609) of garnet mica schist were analysed. These

contain the typical M1 mineral assemblage garnet-phengite-albite-quartz (van der Maar and

Jansen, 1981). IO9609 also contains large chloritoid porphyroblasts up to 5 mm in

length. Zircons from these samples are generally small (< 100 µm) with irregular,

commonly rounded shapes displaying some pitting typical of abrasion during a

sedimentary cycle. The internal zircon structures ranged from regular osillatory zoning,

which appeared to represent only one period of growth, to complex, irregular zoning with

multiple sets of rims overgrowing cores (Figure 3-9). As a reconnaissance study, only

Chapter 3 69eleven zircons from IO9606 and seven zircons from IO9609 were analysed and these

yielded scattered age populations ranging from ca . 1860 to 415 Ma and ca . 1030 to 270

Ma, respectively (Figure 3-7, Figure 3-8).

0.04

0.06

0.08

0.10

0.12

0.14

0 5 10 15 20

1000

500 400 Ma

2000

238 U / 206 Pb

207

Pb /

206

Pb

IO9606n = 11

0

1

2

400 500 600 700 800 900

No.

of

Ana

lyse

s

Age (Ma)

Figure 3-7: Combined histogram with 25 Ma bin widths and kerned probability density curve forzircons sample IO9606. Inset is a Tera-Wasserburg Concordia diagram showing all analyses. Note thatone analysis at 1859 Ma is not included on the histogram or age probability curve.

All of the zircons are interpreted as inherited detrital grains, and the significance of

these results is that the youngest zircon ages in the sample place a maximum age

constraint on the timing of sedimentation of the protolith from which these paragneisses

formed. Thus, for IO9606, the sedimentary protolith can be no older than ca . 415 Ma

(Silurian). For I09609 the three youngest ages have unusually large errors (± 30 Ma, 1σ)

due to the uncertainty on the measurement of the 206Pb/238U ratio. The youngest age ( ca .

270 Ma) is impossibly young (i.e. younger than the orthogneiss that intrudes it),

suggesting there may be some Pb loss (or U gain) affecting some of the zircons in this

sample. The youngest age with better precision (± 10 Ma, 1σ) is at ca . 396 Ma, and this

is considered to be a reliable maximum estimate for the age of the sedimentary protolith

and is consistent with the youngest age from IO9606. That suggests the sediment was

deposited during or after the Devonian. The ca . 330-300 Ma age of the orthogneiss

protolith which intrudes both IO9606 and IO9609 places a minimum age constraint on the

timing of sedimentation, supporting the formation of both paragneiss protoliths between

the Early Devonian and Early Carboniferous.

70 Permo-Carboniferous

0.040.06

0.080.10

0 5 10 15 20 25 30

1000500 400 300 Ma

0.12

238 U / 206 Pb

207

Pb /

206

Pb

0

1

2

0 200 400 600 800 1000 1200Age (Ma)

IO9609n = 10

No.

of

Ana

lyse

s


Figure 3-9: (opposite page) Cathodoluminescence images of zircons from samples identified asCycladic basement. a) Zircon from IO9403 showing regular oscillatory zoning, with an elongate,euhedral morphology with sharp terminations and no core structures. b) Two zircons from IO9607showing more complex internal structures, with an obvious inherited core in the lower zircon. c) IO9609a fractured zircon grain showing a highly luminescent core surrounded by oscillatory zoned zircon withlower luminescence reflecting at least two periods of zircon growth. d) A range of zircons from PA9606showing the general elongate euhedral morphology of grains with sharp terminations, regular oscillatoryzoning and rare inherited cores. e) Elongate, regular oscillatory zoned zircons with no core structures. f)Two zircon grains from SK9601, on the left the grain shows no truncations, indicating only one period ofgrowth, while the grain on the right has a rounded detrital-looking core surrounded by oscillatory-zonedgrowth. g) A range of zircons from NX9485 showing the characteristic elongate, oscillatory-zonedmorphology of grains. h) Zircon grains from NX9637 showing distinct rims with low luminescence,truncating and overgrowing cores with bright luminescence. These rims are related to Miocene partialmelting.

Chapter 3 71

a.

50 µm

IO9403 IO9607

100 µm

b.

d.

200 µm

PA9606

100 µm

SK9601

f.

200 µm

NX9485g.

100 µm

NX9637h.

100 µm

PA9601e.

50 µm

IO9609c.

72 Permo-Carboniferous3.4.2 Paros

The orthogneisses of Paros form part

of the lower group of the Marathi Nappe

(Papanikolaou, 1980) and consist of foliated

granites of two types with similar

compositions except for the presence or

absence of pyroxene. The geochemistry of

basal orthogneisses on Paros suggests that

they are S-type granites and they are

considered to be typical of a volcanic arc or

continental collision zone setting (Engel and

Reischmann, 1997). The age of these units

is undetermined, but comparisons with

other islands suggest they could be any age

between Cainozoic and Palaeozoic, although

some evidence of Alpine age contact

metamorphism by the orthogneisses has

been reported (Paraskevopoulos in Papanikolaou (1980)). The orthogneisses of Paros

have been likened to the Basement of other parts of the Cyclades by several workers (van

der Maar and Jansen, 1983; Gautier et al., 1993; Engel and Reischmann, 1997). Two

samples of orthogneiss were analysed to test this hypothesis and sample locations are

illustrated in (Figure 3-10).

3.4.2.1 Paros Orthogneisses

Nineteen zircons from PA9606 and eight zircons from PA9601 were analysed and

gave ages ranging from ca . 440 to 240 Ma and 450 to 280 Ma, respectively (Figure 3-11,

Figure 3-12). Both gneisses have a quartz-muscovite-albite-biotite assemblage with some

augen of feldspar preserved despite the development of a strong Alpine foliation. As for

the Ios orthogneisses, all zircons are elongate, oscillatory zoned and of magmatic

appearance. Some grains were up to 400 µm in length, and typically showed little

inheritance (Figure 3-9, Figure 3-13). Mixture modelling of dating results (see Appendix

D) identified the main age populations for the samples at 300 ± 1 Ma from thirteen grains

analysed from PA9606 (Table 3-1) with small populations at 332 ± 2 and 297 ± 5 Ma (n

= 4 and 3 respectively) in PA9601. These age groupings suggest that both of the

protoliths to the Paros orthogneisses were emplaced at ca . 300 Ma, close to the time of

emplacement of those on Ios. The interpretation of these ages as emplacement ages

follows the same reasoning as that described in Section 3.4.1.1 for the orthogneisses of

Ios.

0 2 km

N

Upper Unit

OrthogneissMarble

Alluvium

Schist/AmphiboliteGranite

PA9606

PA9601

Figure 3-10: Sample location map for Paros,adapted from Robert (1982) and Papanikolaou(1980).

Chapter 3 73

0.04

0.08

0.12

0.16

0 5 10 15 20 25 30 35 40

200 Ma300400500

238 U / 206 Pb

207

Pb /

206

Pb

PA9606n = 19

0

2

4

6

8

10

200 250 300 350 400 450Age (Ma)

No.

of

Ana

lyse

s

Figure 3-11: Combined histogram with 10 Ma bin widths and kerned probability density curve forzircons from sample PA9606. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

0

0.05

0.10

0.15

5 10 15 20 25 30 35

200 Ma300400

238 U / 206 Pb

207

Pb /

206

Pb

0

1

2

3

250 300 350 400Age (Ma)

No.

of

Ana

lyse

s

PA9601n = 8

Figure 3-12: Combined histogram with 10 Ma bin widths and kerned probability density curve forzircons from sample PA9601. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.


The 440-400 Ma ages reflect inheritance, and no obvious older cores have been

found in this study although 800 and 2260 Ma inheritance has been reported elsewhere

for zircons from Paros orthogneisses (Engel and Reischmann, 1997; Engel and

Reischmann, 1998).

Figure 3-13: Transmitted light photomicrographs of zircons from Basement units, shown in grey-scale.a) Clear, elongate zircon grains with sharp pyramidal terminations typical of magmatic zircon fromPA9606; b) NX9485 similar morphologies to (a); c) NX9320 contains a mixture of two distinct zircontypes; clear euhedral magmatic-appearing grains with sharp terminations and coloured abraded grains withrounded terminations that appear to be detrital from NX9320; d) NX9319 as for (c); e) NX94103 as for (c);f) NX9637 as for (c).

Chapter 3 753.4.3 Sikinos

Basement rocks only crop

out in the southeastern part of

Sikinos and consist of

polymetamorphosed siliclastic

sediments (quartz-chlorite-mica-

garnet schists) intercalated with

subordinate calc-silicate, marble

and meta-acidite. The meta-

acidite is intruded by meta-

diorites and meta-granodiorites

(van der Maar et al., 1981; Franz

et al., 1993). This is slightly

different to the field relationships

on Ios where the schist sequence

is thought to lack marble. One of the meta-diorite bodies was sampled by V. Avdis

(IGME, Athens) and dated in this study. The sample location is illustrated in Figure 3-

14. Franz et al. (1993) suggested that this pre-Alpine basement preserves relics of a pre-

Alpidic amphibolite facies Barrovian metamorphism with peak temperatures of 570-650˚C

and poorly constrained pressures of around 5 kbar.

3.4.3.1 Sikinos Orthogneiss

Van der Maar et al. (1981) suggested that most of the meta-intrusives on Sikinos

have a dioritic composition, with mineral contents and replacement textures very similar to

those found in Ios basement rocks. SK9601 consists of quartz-K-feldspar-plagioclase-

biotite, with zircons mainly hosted by biotite and quartz. Ages from thirty analyses of

twenty-six zircon grains ranged from ca . 860 to 290 Ma (Figure 3-15). Zircon grains

appeared typically magmatic and were consistently around 200 µm in length with regular

oscillatory zoning and only rare core structures. Two main groups of zircons were

identified (Table 3-1); one at 311 ± 1 Ma (n = 12) and the other at 333 ± 1 Ma (n= 12).

The large population at ca . 310 Ma is interpreted as the emplacement age of the

orthogneiss protolith while the older peak represents an early formed inherited component

that may reflect remobilisation of the magma.

0 1 2 km

N

Orthogneiss

Marble

Garnet Mica Schist

Schist

SK9601

Figure 3-14: Sample location map for Sikinos.


0

0.05

0.10

0.15

0.20

0 10 20 30 40

200 Ma3004005001000

238 U / 206 Pb

207

Pb /

206

Pb

0

1

2

3

4

5

6

7

250 300 350 400 450

SK9601n = 29

Age (Ma)

No.

of

Ana

lyse

s

Figure 3-15: Combined histogram with 10 Ma bin widths and kerned probability density curve forzircons from sample SK9601. Inset is a Tera-Wasserburg Concordia diagram showing all analyses. Notethat one analysis at 859 Ma is not included on the histogram or age probability curve.

3.4.4 Naxos

Eight samples of gneiss from the core

of Naxos were analysed to test whether the

core is comprised of pre-Alpine basement

(van der Maar and Jansen, 1983; Andriessen

et al., 1987) or is the high grade equivalent

of the Mesozoic Series rocks which occur

outside the core of the Naxos dome (Jansen,

1973; Jansen and Schuiling, 1976). Sample

locations are indicated in (Figure 3-16). The

so-called Naxos “Basement” consists of

migmatitic gneiss, mica-schists and meta-

pelites which pass with increasing

metamorphic grade and structural depth into

anatectic migmatites and granitic gneisses

(Andriessen et al., 1987). 0 1 2 km

N

Granodiorite

Gneiss/Migmatite

Upper Unit

Ultramafics

SchistMarble

NX9485

NX9314

NX9315 NX9319

NX9320

NX9637

NX9638

NX94103

Figure 3-16: Sample location map of Naxos (adapted from Jansen, 1973; Buick, 1988).

Chapter 3 773.4.4.1 Naxos Layered Acid Gneiss

Two samples (NX9314, NX9485) were taken from the layered acid gneiss (as

described by Buick, 1988) which structurally overlies the core defined by Buick (1988).

This material structurally underlies a tectonic contact marked by ultramafic units (Jansen,

1973; Jansen and Schuiling, 1976; Katzir, 1997) which is used by other workers to

define the Naxos “core” (Andriessen et al., 1987). Because of their location these

samples are possible candidates for being pre-Mesozoic basement. The other samples

were all taken from within the leucogneiss core (as defined by Buick, 1988), with three

samples from the “wispy leucogneiss” lithology (NX9315, NX9319, NX9320) and two

samples from the migmatite core (NX94103, NX9638). Zircons from the two acid

leucogneisses were consistently elongate (> 200 µm in length) and of typical magmatic

appearance, displaying regular oscillatory zoning with little evidence of older core

structures (Figure 3-9). Both NX9314 and NX9485 appear to be meta-I-type granites

consisting of the assemblage quartz-albite-biotite and with relict phenocrysts of

plagioclase and titanite. Twenty-seven analyses of twenty-six zircons from NX9314 and

fifteen analyses of twelve zircons from NX9485 yielded ages ranging from ca . 530 to 270

Ma and 560 to 15 Ma respectively (Figure 3-17, Figure 3-18).

NX9314

0.04

0.05

0.06

0.07

0.08

0.09

0.10

0.11

0.12

5 10 15 20 25

1000

500 400 300 Ma

30

207

Pb /

206

Pb

238 U / 206 Pb

n = 27

0

5

10

15

20

200 250 300 350 400 450 500 550 600

Age (Ma)

No.

of

Ana

lyse

s

Figure 3-17: Combined histogram with 25 Ma bin widths and kerned probability density curve forzircons from sample NX9314. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

The main age populations identified occur at 319 ± 1 Ma (n = 21) for NX9314 and

306 ± 2 Ma (n = 11) for NX9485, with four younger ages in NX9485 identified as

78 Permo-Carboniferoushaving undergone Pb loss (or U gain) from the scatter they produce along the x-axis in

the Tera-Wasserburg diagram in Figure 3-18 (see explanation of interpreting Tera-

Wasserburg diagrams in Appendix D) and because they do not represent clear

overgrowths surrounding the the older zircon populations (Figure 3-9, Figure 3-13). The

one young age of 15 Ma from NX9485 is from a new zircon overgrowth and is

interpreted as being metamorphic in origin. All ages younger than Cretaceous that are

from new zircon growth rims are considered to be metamorphic for reasons which are

discussed in detail in Chapter 6.

NX9485n = 15

0.00

0.05

0.10

0.15

0.20

0 10 20 40

200 Ma300500

100020

7 Pb

/ 20

6 Pb

238 U / 206 Pb30

0

1

2

3

4

5

0 100 200 300 400 500 600

Age (Ma)

No.

of

Ana

lyse

s


3.4.4.2 Naxos Leucogneiss

Samples NX9315, NX9319, NX9320 are all from the southern edge of the

leucogneiss core (as defined by Buick, 1988). They are dominantly quartz-albite-biotite

rocks which are thought to be metamorphosed granitic or feldspar-rich sedimentary rocks

(Jansen and Schuiling, 1976; Buick, 1991; Pe-Piper et al., 1997) and all contain zircons

with similar morphologies and structures. The zircons have irregular shapes, consisting

of elongate, often fractured, regularly oscillatory-zoned grains which are truncated and

overgrown by low luminescent spongy overgrowths of low Th/U metamorphic zircon,

with young growth rims especially concentrated at zircon terminations (Figure 3-9). A

large range of ages is derived from these grains as shown in Figure 3-19, Figure 3-20 and

Figure 3-21. NX9315 contains zircons ranging in age from ca . 205 - 17 Ma (Figure 3-

19). Well-defined zircon age populations occur at 17.5 ± 0.1 Ma (n = 5) and 19.2 ± 0.2

Chapter 3 79Ma (n = 4) and these are discussed in Chapter 6. The abundant 60-10 Ma ages clearly

reflect Alpine metamorphism, appearing as unzoned zircon rims, generally with low

luminescence and low Th/U ratios, surrounding pre-exisiting zircons, features that are

characteristic of metamorphic zircon described in Chapter 6. For this reason, only zircons

having clearly igneous morphologies are used to constrain the time of sedimentation, and

their age represents a maximum age possible for an originally sedimentary precursor. On

this basis, the youngest age that could be considered as a protolith component occurs at

ca . 170 Ma from a clearly oscillatory-zoned grain having a magmatic appearance. Thus

the time at which sample NX9315 was deposited as a sediment is no older than Jurassic

(Figure 3-20).

0

2

4

6

8

10

0 50 100 150 200 250

NX9315

0.02

0.04

0.06

0.08

0.10

0.12

20 Ma304050100200

Age (Ma)

No.

of

Ana

lyse

s

0 100 200 300 400 500

207

Pb /

206

Pb

238 U / 206 Pb

n = 36


Zircons from NX9319 range in age from ca . 320 to 16 Ma. A large Alpine-aged

population of metamorphic zircon overgrowths occurs at 17.7 ± 0.1 Ma (n=16). As in

sample NX9315, the youngest, non-metamorphic, zircon age defined by a regular

oscillatory-zoned grain displaying no growth truncations, occurs at ca . 180 Ma, which

suggests that the sedimentary protolith from which the gneiss was derived is Jurassic or

younger in age.


NX9319

0

0.05

0.10

0.15

0.20

0.25

0.30

0.35

0.40

15 Ma203050100

207

Pb /

206

Pb

238 U / 206 Pb

n = 41

0 100 200 300 400 500 600

0

5

10

15

20

0 50 100 150 200 250 300 350Age (Ma)

No.

of

Ana

lyse

s


NX9320

0

0.05

0.10

0.15

0.20

0.25

0.30

100 50 20 Ma

0 100 200 300 400 500

207

Pb /

206

Pb

238 U / 206 Pb

n = 34

0

1

2

3

4

5

6

7

8

0 50 100 150 200 250 300 350Age (Ma)

No.

of

Ana

lyse

s


Chapter 3 81Similarly, NX9320 shows a wide variety in zircon ages from ca . 340 to 12 Ma

(Figure 3-21). A population of Alpine metamorphic zircon overgrowths occurs at 16.8 ±

0.3 Ma (n = 6). The youngest clearly magmatic-appearing grain constraining the protolith

age occurs at ca . 220 Ma suggesting that the sediment formed in the Triassic or later. It is

interesting to note that there is large population of inherited zircons in this sample aged

322 ± 2 Ma (n = 9) (Table 3-1, Figure 3-13), which is the same age as Basement material

described earlier in this chapter and also found in sample NX9319 which contains zircons

aged 302 ± 4 Ma (n = 3) and 314 ± 3 Ma (n = 4). The Basement age inheritance in these

leucogneiss samples suggests the Basement may have acted as a source for the sediments

now forming the Naxos core.

The existence of detrital zircons of Permo-Carboniferous age, and the presence of

magmatic-appearing Triassic-Jurassic aged zircons in the leucogneiss samples precludes

these gneisses as representing part of the Permo-Carboniferous basement described earlier

in this chapter. The recognition of Cretaceous-aged metamorphic zircon overgrowths

suggests that the sedimentary units were either deposited in the Jurassic and underwent

Cretaceous (Eo-Alpine) metamorphism, or that they were derived from an Eo-Alpine

metamorphic source. In either case the sediments can not be correlated with the pre-

Mesozoic Basement of the Cyclades.

3.4.4.3 Naxos Migmatite

Three samples of migmatite from within Buick’s leucogneiss core of Naxos

(NX94103, NX9638 and NX9637) are thought to represent migmatised granitic or

quartzofeldspathic sedimentary rocks (Jansen and Schuiling, 1976; Buick, 1991; Pe-Piper

et al., 1997). NX9637 is a distinct melt pod within the migmatite. All three samples

consist dominantly of quartz-albite-biotite-muscovite. NX94103 consists of some

notably elongate (up to 400 µm in length) euhedral oscillatory zoned grains (Figure 3-13)

with rare low luminescent new zircon overgrowths on terminations. Other zircons in

NX94103 show similar oscillatory zoning but are shorter and occasionally fractured.

Despite the apparent consistency in zircon morphologies, a range in ages from twenty-six

analyses of twenty-six zircons is found from ca . 350 to 30 Ma (Figure 3-22, Table 3-1).

All of the Cretaceous and younger zircon ages come from fine unzoned overgrowths

surrounding pre-existing zoned cores. The youngest potential protolith age from a clearly

oscillatory zoned grain occurs at ca . 205 Ma, suggesting that the sediments were

deposited in the Jurassic and later metamorphosed in the Cretaceous.


0

1

2

3

0 50 100 150 200 250 300 350 400

NX94103n = 26

Age (Ma)

0

0.05

0.10

0.15

0 50 100 150 200 250

30 Ma4050100200207

Pb /

206

Pb

238 U / 206 Pb

No.

of

Ana

lyse

s


The zircons in NX9637 and NX9638 are very similar to those described from the

wispy leucogneiss units except that NX9638 contains more apparent core structures.

Twenty-one analyses of eighteen zircons from NX9637 reveal a range in ages from ca .

330 to 14 Ma (Figure 3-24) with a large Alpine population of metamorphic overgrowths

at 17.4 ± 0.3 Ma (n = 12) (Table 3-1, Figure 3-9, Figure 3-13). The youngest non-

metamorphic (detrital) zircon population is ca . 225 Ma suggesting that the sedimentary

protolith that melted to form NX9637 is Triassic or younger in age. In NX9638, thirty-

two analyses of twenty-two zircons yields a range in ages from ca . 1870 to 15 Ma (Figure

3-23). The youngest zircon ages which could unquestionably represent the sedimentary

protolith to the migmatite come from two homogeneous oscillatory-zoned grains at ca .

250 Ma, constraining sedimentation to Triassic or younger. As for the wispy leucogneiss

samples, the youngest protolith ages in both NX94103 and NX9638 are Mesozoic,

suggesting that this was a time of sedimentation and that the migmatite, or at least parts of

it, is not derived from pre-Mesozoic basement.

Chapter 3 83

0

1

2

3

4

5

0 100 200 300 400 500

NX9638n = 32

207

Pb /

206

Pb0

0.1

0.2

0 100 200 300 400 500 600

15 Ma203040

238 U / 206 Pb

0.3

Age (Ma)

No.

of

Ana

lyse

s 100

Figure 3-23: Combined histogram with 10 Ma bin widths and kerned probability density curve forzircons from sample NX9638. Inset is a Tera-Wasserburg Concordia diagram showing all analyses. Notethat three analyses at 677, 1059, 1871 Ma are not included on the histogram or age probability curve.

NX9637n = 21

0.00

0.05

0.10

0.15

0.20

0 100 200 300 400 500 600

15 Ma2030

207

Pb /

206

Pb

238 U / 206 Pb

0

2

4

6

8

10

12

14

0 50 100 150 200 250 300 350Age (Ma)

No.

of

Ana

lyse

s

Figure 3-24: Combined histogram with 10 Ma bin widths and kerned probability density curve forzircons from sample NX9637. Inset is aTera-Wasserburg Concordia diagram showing all analyses.


Table 3-1: Summary of U-Pb zircon ages for Basement samples

Sample No.spots

No.zircons

Rock-type Main Ages(No. Analyses)

Age Range

IO9403 26 24 orthogneiss 307 ± 1 (15) 292 - 1090IO9404 8 8 orthogneiss 310 ± 1 (4) 306 - 80589640 13 12 orthogneiss 324 ± 2 (8) 301 - 417IO9607 13 11 leucogneiss 321 ± 4 (3) 296 - 2437IO9606 11 11 garnet mica schist 413 - 1859IO9609 10 7 garnet mica schist 270 - 1028PA9606 19 19 orthogneiss 300 ± 1 (13) 237 - 440PA9601 8 7 orthogneiss 332 ± 2 (4) 284 - 448SK9601 29 25 orthogneiss 311 ± 1 (12)

333 ± 1 (12)287 - 859

NX9314 27 26 layered acid gneiss 319 ± 1 (21) 272 - 533NX9485 15 12 layered acid gneiss 306 ± 2 (11) 15 - 555NX9315 36 30 leucogneiss 17.5 ± 0.1 (5)

19.2 ± 0.2 (4)17 - 205

NX9319 41 29 leucogneiss 17.7 ± 0.1 (16) 15 - 320NX9320 34 28 leucogneiss 322 ± 2 (9)

16.8 ± 0.3 (6)15 - 339

NX94103 26 26 migmatite 28 - 352NX9638 32 22 migmatite 15 - 1871NX9637 21 18 melt pod 17.4 ± 0.3 (12) 14 - 333

3.5 Combined Permo-Carboniferous Results

A combination of all Permo-Carboniferous ages from both primary and detrital

zircons in the Cyclades shows that the majority of zircon ages fall in the range 330-300

Ma (Figure 3-5). Most of these ages are either from Basement rocks in the Cyclades or

from Mesozoic sediments that preserve a Permo-Carboniferous component of zircons as

inherited/detrital grains. The Cycladic Basement records the timing of extensive granitoid

intrusion in the area. The sedimentary rocks that these granitoids intrude contain zircons

as young as 350 Ma old, so we can safely infer that the granites are intruding material

which is no older than Devonian (400 Ma). The intrusion of granites to form basement in

the Aegean region coincided with extensive granite formation during the Variscan

orogeny, and these Carboniferous ages correspond closely to some of the oldest ages

found in rocks from the Pelagonian zone in Greece (Yarwood and Aftalion, 1976;

Mountrakis, 1984).

Chapter 3 85

0

5

10

15

20

25

30

35

240 260 280 300 320 340 360Age (Ma)

No.

of

Ana

lyse

s

Figure 3-25: Density Distribution plot of all Permo-Carboniferous results (overlain on histogram with5 Ma bin widths).

3.6 Discussion

3.6.1 Confirmation of the Existence of Pre-Mesozoic basement

The suggestion that pre-Mesozoic basement exists in the Cyclades (e.g., van der

Maar and Jansen, 1983) has been confirmed by SHRIMP U-Pb zircon dating of

orthogneisses from four islands: Ios, Paros, Sikinos and Naxos. The restricted range in

ages from 330-300 Ma determined for the intrusion of the basement orthogneiss

protoliths, conflicts with previous interpretations of a Pan-African basement age (Henjes-

Kunst and Kreuzer, 1982; Andriessen et al., 1987), although some inheritance of Pan-

African zircons has been identified in this study. The interpretation that the ages represent

the timing of cooling and emplacement of granitoid magmas is supported by the

morphology of the zircons analysed (as described in Section 3.4.1.1), which show no

evidence of new metamorphic growth or recrystallisation, as has previously been

suggested (Henjes-Kunst and Kreuzer, 1982). Although the granitoid precursors to these

gneisses were most likely S-type, there is no evidence to suggest that the zircon ages are

entirely inherited (i.e. that no new zircon growth occurred during granitoid production).

The results of the present study are consistent with the results of two recent studies

utilising the Pb evaporation technique (Kober, 1986) on zircons from orthogneisses on

Naxos and Paros. These confirm that the intrusions forming the basement on these

islands are no older than Variscan in age. Reischmann (1997; 1998), reports an age of

86 Permo-Carboniferousca . 280 Ma from the migmatitic core of Naxos whereas Engel and Reischmann (1997;

1998) report an age of ca . 310 Ma for the basal orthogneisses of Paros. These results

confirm that the granites forming the Basement of the Cyclades are Variscan in age.

3.6.2 The Timing of Pre-Alpine Metamorphism M0

The timing of amphibolite facies metamorphism (M0) is bracketed by the intrusion

of orthogneiss precursors and Alpine M1 metamorphism. Thus M0 must post-date the

time of intrusion of the orthogneiss precursors at ca . 330-300 Ma, and pre-date the Alpine

high-P M1 metamorphism at ca . 50 Ma (see Chapter 6). As the Cycladic basement

preserves evidence of structural fabrics thought to pre-date the Alpine orogeny (e.g.,

Grütter, 1993) the formation of these structures must also be constrained to the time

period 330-50 Ma. However, pre-Alpine granitoids in the Menderes Massif are thought

to have been converted to orthogneisses during the Alpine orogeny (Hetzel and

Reischmann, 1996) and it is possible that the early fabrics recorded by Cycladic basement

rocks and not found in Mesozoic Series rocks reflect the potentially different Alpine

tectonic histories of these units, and do not result from a pre-Alpine event. This

possibility will be hard to test because the timing of juxtaposition of the Basement and

Series rocks remains unclear.

3.6.3 Complications within the Naxos core

Eight samples from the core of Naxos, as defined by (Jansen and Schuiling, 1976)

according to the distribution of ultramafic lenses, were dated to confirm whether the core

was entirely composed of pre-Mesozoic basement (Andriessen et al., 1987). The results

suggest that although a restricted unit of orthogneiss outside the zone of partial melting

(the layered acid gneiss of Buick, 1988) is pre-Mesozoic basement, the rest of the core is

probably derived from Mesozoic sediments. This confirms the view of Jansen and

Schuiling (1976) that the core is the same age as the Mesozoic Series rocks surrounding it

(see Chapter 4 for a discussion of the Mesozoic history of Series metasediments). The

existence of at least some Series rocks, in the form of marble and pelite “rafts” within the

partially melted core of Naxos, has previously been recognised (Jansen, 1973; Jansen and

Schuiling, 1976). These rafts appear to “float” within a quartzofeldspathic leucogneiss

which, with progressive partial melting, develops into a migmatite while the rafts remain

essentially unmelted (Buick, 1988; Buick, 1991; Buick and Holland, 1991). The rafts

show remarkable similarities to the Series rocks outside the core (Buick, 1988), however

the enveloping leucogneiss has no obvious correlative outside the core (Buick, pers

comm.)

Zircons from the leucogneiss and migmatite of the Naxos core display a spread in

ages and a range of morphologies indicative of sedimentary host rocks (Figure 3-9). The

zircons are often rounded and occasionally overgrown by thin (< 30 µm) new

Chapter 3 87metamorphic zircon growths ranging in age from Cretaceous to Tertiary. The youngest of

these records the timing of partial melting (see Chapters 5 & 6). Though the age range for

the rocks spans 12 - 1871 Ma, ages which could be considered as representative of the

protolith ages, as opposed to a metamorphic age, are mainly restricted to the Triassic-

Jurassic, with the youngest protolith ages (i.e. those constraining the timing of

sedimentation) being Jurassic. The identification of these sedimentary zircons in all of the

six leucogneiss and migmatite samples suggests that the majority of the leucogneiss core

is derived from a sedimentary protolith of heterogeneous age containing some Triassic-

Jurassic material, and contradicts the identification of a “Central Granite” in the gneissic

core of Naxos. The supposed preservation of igneous textures from this material (Pe-

Piper et al., 1997), is equally well-explained as the result of partial melting which has

affected the entire leucogneiss core (Buick, 1988). The chemical arguments suggesting

that the leucogneiss core is a granite (Pe-Piper et al., 1997) are also inadequate as it has

been demonstrated that partial melting can produce chemical signatures that mimic those

of S-type granites (Watt and Harley, 1993). The possibility that these zircons are purely

inherited can only occur in S-type granites which are undersaturated with respect to

zircon, in which case dissolution of zircon grains would occur (Watson and Harrison,

1983). A zircon population dominated by inheritance would also be expected in a granite

containing a large quantity of restite (see Keay et al. paper in Appendix A), where low

volumes of melt would restrict new zircon growth. In this scenario, large volumes of

unmelted sedimentary rock should be identifiable in the Naxos core but field observations

do not support this hypothesis because, as Pe-Piper et al. (1997) point out, the core

dominantly preserves igneous textures. The Central Granite of Pe-Piper et al. (1997)

would have to predate melting within the leucogneiss core, which is well-constrained by

dating of a migmatitic melt pod (NX9637) in the migmatite at 17 Ma (Table 3-1). All of

the new metamorphic zircon growth in other samples from the Naxos core are the same

age or younger than NX9637, suggesting that new zircon was produced during the partial

melting episode associated with M2. If the core does represent a “Central Granite” as

proposed by Pe-Piper et al. (1997), it seems peculiar that no record of it is preserved in

the included zircons. The Mesozoic history of the Cycladic rocks is discussed in more

detail in the following chapters.

3.6.4 Correlations with North Africa

According to most plate reconstructions for the area, the Cyclades were attached to

the northern margin of Gondwana during the Carboniferous (Robertson and Dixon,

1984). For this reason, Late Carboniferous granitoids associated with the Variscan

orogeny should be abundant in North Africa. This is certainly the case in west and

northwest Africa, which have close ties with the Alpine orogenic belt (Cahen et al., 1984;

88 Permo-CarboniferousDallmeyer and Lecorche, 1991), but insufficient data exist for north Africa to comment on

the abundance, or lack thereof, of Late Carboniferous intrusions.

3.6.5 Correlations with the Menderes Massif, Turkey

The basal units of the Cyclades and the Menderes Massif both consist of deformed

orthogneisses intruding a schist mantle and many workers have suggested these units are

homologous (Dürr et al., 1978; Jacobshagen et al., 1978; Dürr, 1986; Okay, 1989).

However, Late Carboniferous magmatic rocks are scarce in Turkey (Sengör et al., 1991)

and the age of the Menderes Massif orthogneiss protoliths have been dated consistently

as latest Proterozoic, ca . 550 Ma (Kröner and Sengör, 1990; Loos and Reischmann,

1995; Hetzel and Reischmann, 1996; Dannat and Reischmann, 1997) using the Pb-Pb

single zircon evaporation technique. Intrusive rocks associated with this period at the

Cambrian-Precambrian boundary have been interpreted as the products of a magmatic

episode contemporaneous with, or slightly younger than, the end of the Pan-African

episode (Loos and Reischmann, 1995). If the observation of the intrusive contact

between schist and augen gneiss is correct (Erdogan, 1992) then the gneiss protoliths

must be intruding Precambrian schists. No rocks of this age have been found in the

Cyclades, although they may have contributed to the 550 Ma age peak found as

provenance ages described in Chapter 2. It was also noted in Chapter 2 that there are

distinct differences in the Early Palaeozoic histories of the Cyclades and the Menderes

Massif, which further suggests that they have undergone a different early evolution.

While no correlation between the basement of the Menderes Massif and the

Cyclades seems possible, there are similarities between the Basement of the Cyclades and

that of the Turkish Pontides (Figure 1-5). The Bayburt area of the Pontides has a

polymetamorphosed basement comprised of schists intruded by Late Carboniferous

granites (Cogulu (1975), in German, quoted in (Sengör et al., 1980), which is overlain

by a sequence of Permo-Triassic marine sediments. This area, along with the Istanbul

and Kirklareli “nappes” of western Turkey, has been correlated with the Pelagonian zone

of mainland Greece (Mountrakis, 1986), which also shows many similarities to the

Cyclades (as is discussed in the next section).

3.6.6 Correlations with the Pelagonian Zone, Internal Hellenides, Greece

Both the Pelagonian zone of mainland Greece and the Cyclades are thought to have

formed on a thick Variscan basement and similarities in the geology of the two areas have

often been noted (Jacobshagen et al., 1978; Blake et al., 1981; Robertson and Dixon,

1984; Jacobshagen, 1986). Unlike the Menderes Massif, there is evidence of Late

Carboniferous granitoid intrusion in the Pelagonian zone, with U-Pb dating of zircon

from the Kataphygion granodiorite yielding an age of ca . 300 Ma (Yarwood and Aftalion,

1976) and a similar U-Pb age is reported for the Kastoria granites (Mountrakis, 1984).

Chapter 3 8940Ar-39Ar dating of minerals from granites in the Olympos region also constrain the time

of intrusion of these magmas to ca . 300 Ma (Schermer et al., 1990; Lips et al., 1997).

These ages suggest that the basement of the Pelagonian zone and the Cyclades may be

correlated, both showing the influence of Late Carboniferous granitoid intrusions during

the late stages of the Variscan orogeny. The Pelagonian granitoids have a subduction-

related component thought to be the product of subuction along the northwest margin of

PaleoTethys (Pe-Piper et al., 1993).

The existence of Mesozoic quartzo-feldspathic sediments in the Naxos core may be

correlated to similar rocks identified in the Pelagonian zone. Mountrakis (1984) identifies

an Upper Palaeozoic (Permian to early Triassic) sedimentary sequence in the Pelagonian

zone that consists mainly of meta-arkoses, meta-pelites, phyllites, phengite schists, meta-

sandstones, quartzose conglomerates, tuffs, limestones and calc-silicates. The meta-

arkoses are mainly composed of detrital K-feldspar and were deposited in close proximity

to a granite source. Further discussion of correlations between the Mesozoic rocks of the

Cyclades and the Pelagonian zone is left to Chapter 4.

3.6.7 Correlations with the External Hellenides, Crete

Similarities in the age and geologic character of the basement to the Cyclades and

amphibolites, gneisses and mica schists of the External Hellenides, exposed as the

Tripolitza unit on Crete, have been noted (Seidel et al., 1982; Bonneau, 1984). In

addition, the basement on Crete records evidence of a possible Variscan M0

metamorphism of Barrovian-type amphibolite grade (Seidel, 1977). Alpine M1 high-P

metamorphism is recorded by the development of blue amphibole (riebeckite, crossite),

but the formation of albite, chlorite and epidote are attributed to an earlier metamorphic

episode (Seidel, 1977). K-Ar dating of this basement yields muscovite ages spanning

100 Ma (315-205 Ma), whereas ages from hornblende are more restricted (300-270 Ma).

The hornblende ages might be recording the timing of M0 immediately following the

intrusion of the Variscan granitoids in the Cyclades. The preservation of Variscan ages

has been used to suggest that the basement was only weakly affected by M1. This

interpretation is supported by the typical correlation of decreasing K-Ar dates with

decreasing Niggli-Mg values of the hornblendes predicted by degassing experiments in

vacuum (O'Nions et al., 1969). The prevalence of Late Carboniferous ages for the

basement of Crete supports its correlation with that of the Cyclades.

3.6.8 Tectonic Implications of Age data

The identification of voluminous Late Carboniferous granitoid intrusions in the

basement rocks of the Cyclades, and their correlation with basement in the Pelagonian and

the External Hellenides implies that a large area of the southern portion of the Alpine chain

was influenced by the Variscan orogeny prior to Alpine events. The lack of granitoids in

90 Permo-Carboniferousthis age range in the Menderes Massif suggests the Cyclades may represent the south-

eastern extent of the influence of the Variscan orogeny. This could have important

implications for tectonic reconstructions of the area, especially in relation to the position

of these areas on the margin of North Africa prior to their separation as microcontinental

blocks in the Jurassic (Robertson and Dixon, 1984).

3.7 Synthesis

SHRIMP U-Pb zircon ages ranging between 330-300 Ma demonstrate that the

orthogneissic Basement of the Cyclades is comprised of Late Carboniferous intrusives

emplaced during the late stages of the Variscan orogeny. Provenance ages from the

garnet-mica schists intruded by these granitoids indicate a maximum depositional age of

Early Devonian. Unless there is older basement that remains unexposed, the Cyclades

consist of distinctly younger crust than the Menderes Massif, where orthogneiss

protoliths are believed to be ca . 550 Ma in age. There are some silimarities between the

Cyclades and the Pontides however, as discussed in Section 3.6.5. The Cycladic

Basement also shows similarities to the basement of the Pelagonian zone of mainland

Greece and to the External Hellenides exposed on Crete where Late Carboniferous

granites have also been reported. The identification of Late Carboniferous granitoids

intruding post-Devonian sediments in the Cyclades confirms that the basement is Variscan

in age and conflicts with earlier interpretations of this period as one of high grade M0

amphibolite metamorphism resulting in the recrystallisation of the orthogneisses (Henjes-

Kunst and Kreuzer, 1982). Zircons from these rocks show no evidence of having

undergone recrystallisation. The dominant ca . 300 Ma age population found in

orthogneisses is from generally well-formed, elongate, euhedral zircons with

characteristically igneous morphologies. A pre-Alpine amphibolite-facies metamorphism

(M0) does affect the basement of the Cyclades, but this must post-date the Variscan

magmatism at 300 Ma and pre-date Alpine metamorphism (M1) at 50 Ma. Correlations

with Crete suggest M0 might occur at ca . 270 Ma. The Series rocks which structurally

overlie the Cycladic Basement do not retain any evidence of a “pre-Alpine” metamorphic

and structural history. Dependent on when the Series and Basement rocks were

juxtaposed (which remains poorly constrained), this could also be taken as evidence to

suggest M0 was pre-Mesozoic.

Chapter 4 91

4. TRIASSIC / JURASSIC GEOLOGICAL EVOLUTION OF

THE CYCLADES

4.1 Introduction

The depositional age and character of the Cycladic Series rocks are the focus of this

chapter. The ages were constrained by identifying and dating the youngest detrital zircon

component in these rocks. New age constraints on the crystallisation of Syros’s Vari

orthogneiss, a unit of previously controversial age, show that it is Triassic in age. From

Permian to Triassic times, a major change in the shape of the supercontinent Pangea

occurred, with substantial east-west movement of Eurasia and North America relative to

Africa. This restructuring is evidenced by extensive rifting and widespread Triassic

volcanism in the Alpine region (Hynes, 1974; Celet et al., 1977), marking the early stages

of the break-up of Pangea (Gealey, 1988). It was during this time that slivers of

continental crust first began to rift away from the northern margin of Gondwana and

move northwards toward Eurasia (Robertson and Dixon, 1984). In the Mediterranean

area, evidence of rifting and associated volcanism can be found in the Dinarides of former

Yugoslavia (Pamic, 1984) and the internal and external Hellenides in Greece (Seidel et

al., 1981; Pe-Piper, 1982; Harbury and Hall, 1988; Magganas et al., 1997) and

southwest Turkey (Robertson and Woodcock, 1981). Basin formation and deposition of

shallow marine sequences occurred throughout the Mesozoic. Following rifting,

volcanism and sedimentation, collision of the African and Eurasian plates caused high-P

metamorphism of the rock units associated with the Alpine orogeny (Chapter 6). As a

result, the Series rocks of the Cyclades contain some of the most famous blueschist

localities in the world (Okay, 1989). The age of the blueschist protoliths are broadly

termed “Mesozoic” because there are very few age constraints available for these units

other than metamorphic ages. Sparse palaeontological identifications range from Triassic

to Cretaceous in age, with most units considered to be Triassic (Dürr et al., 1978;

Bonneau, 1984). Interleaved bauxite units within Series marbles are thought to be

Jurassic in age on the basis of geochemical similarities with bauxites of this age in eastern

Europe (Feenstra, 1985).

92 Triassic/Jurassic

Equator

LATE TRIASSIC

Africa

Antarctica Australia

India

SouthAmerica

NorthAmerica

CimmeriaChina

~220 Ma

Pantha

lassa

Pantha

lassa

PA

NG

EA

Tethys

Cyclades

Figure 4-1: Late Triassic plate reconstruction showing the active margin north of Africa and thelocation of the Cyclades and Tethys. The initial stages of rifting in this region began when theCimmerian continent moved northward to collide with the southern margin of Eurasia (Reconstructionsfrom Dercourt et al., 1986; Ricou, 1994; Metcalfe, 1996).


Overlying the basement rocks described in Chapter 3 is a sequence of interbedded

marbles, metavolcanics and metasediments that form the Series rocks of the Cyclades.

Deposition of Series sediments is generally thought to have occurred in the shallow

marine environment of a continental margin (e.g., Sifnos Okrusch et al., 1978). The

distribution of Series rocks in the Cyclades is illustrated in Chapter 1. Along with the

Basement, the Series rocks have also undergone the same history of M1 Eocene high-P

low-T metamorphism and M2 Oligo-Miocene greenschist to amphibolite facies

metamorphism prior to Mio-Pliocene exhumation. The nature of the contact between the

Basement and Series rocks is poorly constrained. Early work suggested that the Series

sediments were deposited directly onto the Variscan basement (Dürr et al., 1978), but

other workers have suggested a tectonic contact between the two units (e.g., Paros

Papanikolaou, 1980). Although certain workers have interpreted the the contact between

Series rocks and the Cycladic Basement as a thrust fault on Ios (van der Maar and Jansen,

1983; Grütter, 1993) and Sikinos (van der Maar et al., 1981; Franz et al., 1993), others

have interpreted the contact as a normal (detachment) fault (Lister et al., 1984). On

Naxos the contact between the two units is delineated by a discontinuous horizon of

ultramafic rocks. There have been two contrasting interpretations for the origin of this

ultramafic layer as either remnants of oceanic crust, emplaced along a major tectonic

contact inferred to be a thrustplane (Jansen and Schuiling, 1976), or Katzir (1997)

hypothesises that the ultramafics are relict peridotites which have been incorporated into

the Naxos leucogneiss prior to M2 metamorphism during the Eocene (M1) collisional

process where the underthrusting continental slab (leucogneiss) scavenged part of the

Chapter 4 93overriding subcontinental mantle. As both the Basement and Series have undergone the

same Tertiary metamorphic history, it is assumed they were juxtaposed prior to this time,

although an alternative suggestion is that they were juxtaposed during the Alpine M1

metamorphism (Katzir, 1997).

Interbedded with marbles of the Series rocks on Naxos are emery deposits

representing original karst bauxites formed during a period when the carbonate platform,

from which the marble sequences were derived, was emergent (Feenstra, 1985).

Bauxites are also preserved in the Series rocks on the islands of Ios, Sikinos, Paros and

Iraklia (Dürr et al., 1978). Meta-bauxites and karst bauxites are distributed throughout

the Menderes Massif of Turkey and the Pelagonian zone of Greece (Figure 3.2b). These

deposits are not recognised in the northern islands of the Cyclades (Sifnos, Syros, Tinos,

Andros) leading (Blake, 1980) to speculate that there are lithological differences between

the north and south Cyclades. No Basement is recognised in the northern Cyclades,

although Syros contains one orthogneiss unit of uncertain age, the Vari gneiss, which

was investigated during this study.

The Vari unit of Syros consists of an augengneiss mantled by metasediments and is

considered by most workers to be allochthonous (Seidel et al., 1976; Bonneau, 1984).

The Vari gneiss is a greenschist-facies quartzofeldspathic rock containing relicts of

epidote-amphibolite assemblages related to an earlier metamorphic event of unknown age.

The exact relationship between this unit and the surrounding metasediments is unclear.

Bonneau et al. (1980) and Blake et al. (1981) originally interpreted the Vari unit as a

window under the Syros schists (which are equivalent to the Series rocks of other

Cycladic islands). Maluski et al. (1987) suggested that the unit was overthrust onto

Series rocks while Ridley (1984) suggested that the unit overlies a listric normal fault and

is from higher in the synmetamorphic structural pile than the Series rocks, which suggests

it forms part of the Cycladic Upper Unit. The gneiss is considered by some to be

equivalent to the Asteroussia nappe which forms the highest structural level on Crete and

is considered to be equivalent to the Upper Unit of the Cyclades (Seidel et al., 1976;

Seidel et al., 1981; Bonneau, 1984). Hecht (1984) suggested that the unit is Early Alpine

or Variscan in age and has undergone Eocene high-P metamorphism M1, while Ridley

(1984) suggested that the Vari gneiss has undergone a lower pressure paragenesis than

the Series rocks of Syros. No relict M1 blueschist assemblages have been reported from

the Vari gneiss, but these can be difficult to identify in orthogneisses because of their

mineralogies, as is the case on Ios (van der Maar and Jansen, 1983; Grütter, 1993).


Although potentially dateable metavolcanic units are interbedded with marble in the

Cyclades, no protolith ages have been reported from these rocks. A variety of ages from

94 Triassic/JurassicAlpine metamorphic minerals contained within the Series rocks have been reported and

these will be discussed in Chapter 6. The age of the sediments which comprise the Series

rocks of the Cyclades has previously been constrained by sparse palaeontological

identifications. Rare fossils of the algae Triploporella cf. remesi (Steinmann) constrain

the age of some low grade Series rocks containing dolomite on Agrilos (south of Naxos)

to be Late Jurassic to Early Cretaceous (Dürr et al., 1978). According to Maluski et al.

(1987), early Mesozoic fossils occur in rocks from Sifnos and Kynthos, while dolomites

on Tinos contain Late Triassic algae and coral fragments (Melidonis, 1980). A

carbonaceous calcite marble from the Series rocks of Ios was identified as

STYLOSMILIA , typical of the Mid-Jurassic to Early Cretaceous (Grütter, 1993). These

constraints suggest that the age of the Series rocks spans most of the Mesozoic. Other

age constraints have come from lithological correlations and isotopic constraints. The

emery deposits of Naxos have been correlated with unmetamorphosed bauxite deposits of

Jurassic age in Greece and Yugoslavia on geochemical grounds (Feenstra, 1985). An

Early Mesozoic age for some sediments on Naxos is supported by an initial 87Sr/86Sr

ratio of 0.7080 ± 0.0001 for Naxian marbles (Andriessen et al., 1979) thought to be

representative of the original marine limestone protolith ratio which approaches the known

isotopic composition of marine strontium in the Early Mesozoic (Veizer and Compston,

1974).

The problem of the time of formation of the Vari gneiss, Syros, can be related to its

mineralogy, previous radiometric dating results and conflicting field observations.

Maluski (1987) reports a 40Ar-39Ar plateau age of 75 ± 3 Ma from phengite from the

gneiss, and although this age is from disturbed argon spectra, it was considered to give

the age of a major tectonometamorphic event. For this reason the Vari unit has been

likened to the Upper Unit of the Cyclades that also records Cretaceous ages for

metamorphism of this material e.g., on Anaphi (Reinecke et al., 1982), Nikouria,

Amorgos, Donoussa (Dürr et al., 1978) and Tinos (Patzak et al., 1994) as well as south

of the Cyclades in Crete (Seidel et al., 1981) and northwest in Evvia (Maluski et al.,

1981). Some of these ages were interpreted as recording the timing of a metamorphic

event, while others (e.g., Anafi and Tinos) were considered to represent both the timing

of metamorphism and ophiolite formation. However, these units consist of distinctly

different lithologies to the Vari gneiss, being mainly medium to low grade metasediments

and ophiolites rather than granites. The preservation of Cretaceous ages in phengites

from the Vari gneiss suggests that the unit did not undergo the Eocene high-P

metamorphism (M1) which would have been expected to reset the argon isotope

systematics. However, if Maluski’s argon ages are the result of an excess argon

component, as has been found in other argon studies of the Cyclades (e.g., Altherr et al.,

1979; Grütter, 1993), the above argument does not apply.

Chapter 4 95Maluski et al. (1987) also report one 40Ar-39Ar phengite age of 30.3 ± 0.9 Ma from

a rock in contact with the Vari unit. Ages of 48 Ma obtained at low steps from this

sample are thought to indicate that the 30 Ma age reflects a secondary opening of the

system which was related to tectonic emplacement of the Vari nappe on top of the Syros

schists, and is used as further support to the argument that the Vari unit was tectonically

emplaced after M1 (Maluski et al., 1987). Recent field work by Freiburg and Tuebingen

universities has led to a completely different interpretation however, based on the

identification of an apparent intrusive contact between the Vari gneiss and the surrounding

Syros schists (C. Ballhaus, pers comm). This suggests that the Vari gneiss protolith

could be much older than previously supposed, because the schists have demonstrably

experienced Eocene high-P metamorphism and are thought to be Mesozoic in age,

forming part of the Cycladic Series rocks. Both the age of the Vari Gneiss and some of

the Syros schists is investigated in this chapter to resolve many of these issues.

4.4 SHRIMP U-Pb Results for the Vari Gneiss

Sample SY9603 is from the Vari

gneiss, a strongly foliated orthogneiss

located in the southeast corner of the island

of Syros (Figure 4-2) consisting of quartz-

feldspar-muscovite-epidote. Brief

descriptions of the zircon morphologies

present in the Vari gniess sample are listed

in Appendix E. Zircon grains are generally

clear, less than 200 µm long and have

slightly rounded morphologies. CL

imaging revealed that most zircons display

oscillatory zoning and there are no obvious

core structures indicating the presence of

inherited zircon grains (Figure 4-7). A

small amount of apparently new zircon

growth was identifiable close to grain

terminations, but analysis of these zones

yielded similar ages to the more regular

oscillatory zoned parts of the zircon,

suggesting that the growths were formed at the same time.

Analyses of twenty-five zircons from the Vari unit yielded a cluster of Triassic ages

depicted in Figure 4-3 (and listed in Table 4-1).

0 1 2 km

N

Metabasite

Orthogneiss

Marble

Schist

SY9603

89646

SY9630

Figure 4-2: Island of Syros showing samplelocation.


0

0.05

0.10

0.15

0.20

10 20 30 40

200 Ma300

Common Pb

SY9603n = 25

207

Pb /

206

Pb

238 U / 206 Pb

0

2

4

6

8

10

12

200 220 240 260 280 300

No.

of

Ana

lyse

s

Age (Ma)

Figure 4-3: Combined histogram with 10 Ma bin widths and kerned probability density curve forzircons from sample SY9603. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

Analysis of twenty-five zircon grains produced a spectrum of ages ranging from ca .

255 to 210 Ma. Four analyses yielded anomalously young ages and plot to the right of

the main group of analyses on Concordia, consistent with these analyses being affected

by Pb loss. There was no obvious correlation between Pb loss and the type of grain

analysed or uranium content. Neglecting these four ages and processing the other results

through a mixture modelling procedure yielded an age of 240 ± 1 Ma (1σ) from twenty-

one zircons. This result is interpreted to be the age of magmatic crystallisation of the Vari

gneiss protolith. None of the twenty-five zircons analysed show any evidence of

inheritance even though the analytical strategy targetted all of the different internal zircon

structures (Figure 4-7).

4.5 Depositional Age and Provenance of Sediments in Cyclades

Zircons from eighteen samples of Mesozoic Series rocks from six of the Cycladic

islands were analysed and the youngest detrital zircon age was used to constrain the time

at which these units were deposited. The resultant provenance ages are mainly Jurassic-

Triassic in age, and these grains were often overgrown by younger metamorphic

overgrowths which ranged in age from Cretaceous to Tertiary, e.g. the Naxos

Leucogneiss samples discussed in Chapter 3.

Chapter 4 974.5.1 Syros

Two samples from Syros were analysed, and their location is indicated in Figure 4-

2. 89646 is a quartzite (also analysed by Baldwin, 1996) while SY9630 is a foliated,

quartz-feldspar-muscovite schist. Both samples form part of the “Schists of Syros”

(Hecht, 1984). Sample 89646 contains an assortment of small (~ 100 µm in length)

whole and broken zircons, many with surface pitting indicative of abrasion, which would

be indicative of them having experienced a prolonged sedimentary cycle. Some zircons

contain complex core structures. Many of the grains are oscillatory-zoned (Figure 4-7),

but growth zoning is commonly truncated by later overgrowths and most grains appear to

record multi-stage histories.

0

0.05

0.10

0.15

0.20

0.25

0.30

0.35

0 20 40 60 80 100

75 Ma100200300500

89646n = 27

207

Pb /

206

Pb

238 U / 206 Pb

0

1

2

3

4

5

6

7

0 100 200 300 400 500 600 700

No.

of

Ana

lyse

s

Age (Ma)

Figure 4-4: Combined histogram with 20 Ma bin widths and kerned probability density curvefor zircons from sample 89646. Inset is Tera-Wasserburg Concordia diagram showing all analyses.

Analyses of twenty-seven zircons from sample 89646 reveal ages ranging from ca .

650 to 75 Ma (Figure 4-4). Six ages are Cretaceous (the youngest in the population) and

come from unusual irregular or sector-zoned grains or from homogenous rims

surrounding oscillatory-zoned cores. The youngest protolithic age from detrital grains

occurs at ca . 225 Ma and is from a clear, euhedral oscillatory zoned zircon. 89646 also

has a large population of zircons at 288 ± 2 (7) suggesting a Permo-Carboniferous source

was important in providing material for the formation of the quartzite.

Twenty-one zircons were selected from sample SY9630 for analysis. In contrast to

sample 89646, the zircons from SY9630 were generally larger (> 200 µm in length),

98 Triassic/Jurassicmore angular, showed little evidence of abrasion, and exhibited regular oscillatory growth

zoning. This sample also showed a smaller range in ages from ca . 270 to 107 Ma (Figure

4-5). The six youngest ages are Late Jurassic/Cretaceous forming two small populations

at 108 ± 3 Ma (2) and 135 ± 2 Ma (4). All come from oscillatory zoned grains that show

little evidence of truncations in growth, unlike the Cretaceous rims described from 89646.

This suggests that the Cretaceous ages are actually protolith ages rather than overgrowths,

and that the sample is most likely derived from a Cretaceous or younger sedimentary

protolith. While SY9630 has some Jurassic-Triassic ages at 199 ± 1 Ma (6) and 221 ± 1

Ma (9).

0

1

2

3

4

5

6

100 150 200 250 300 350

No.

of

Ana

lyse

s

Age (Ma)

SY9630n = 23

0

0.05

0.1

0 10 20 30 40 50 60 70

100 Ma200300

207

Pb /

206

Pb

238 U / 206 Pb

Figure 4-5: Combined histogram with 10 Ma bin widths and kerned probability density curve forzircons from sample SY9630. Inset is aTera-Wasserburg Concordia diagram showing all analyses.

Chapter 4 99

4.5.2 Naxos

Ten samples of the Series rocks on

Naxos were analysed and their locations are

illustrated in Figure 4-6. These included;

two samples of quartzite (NX9481,

NX9451), six samples of calc-silicate

(NX9461, NX9463, NX94112, NX9464,

NX94120, NX94121) and two samples of

pelite (NX9490, NX94106). Both NX9451

and NX9481 are quartz-dominated rocks

containing feldspar and epidote. In addition,

NX9451 contains chlorite whereas NX9481

contains biotite. NX9451 is a low grade

metasediment from the chlorite-sericite zone

defined by (Jansen, 1973). It is in close

proximity to the biotite isograd and estimated

to have attained a temperature of ~ 500 ˚C

(Jansen and Schuiling, 1976). NX9481 is

from the sillimanite zone (Jansen, 1973)

where temperatures are thought to have

reached 620-660 ˚C (Jansen and Schuiling, 1976).

Both rocks appear to have been recrystallised during deformation, with quartz

forming ribbons of polygonal subgrains. Zircons are mainly included in quartz and

feldspar and are generally between 100 - 200 µm in length, with regular oscillatory

zoning. They contain only rare zircon core structures and show little evidence of new

metamorphic zircon growth (Figure 4-7). There are no signs of surface pitting or

abrasion, although some grains have irregular outlines due to resorption.

0 1 2 km

N

Granodiorite

Gneiss/Migmatite

Upper Unit

Ultramafics

SchistMarble

NX9451

NX9481

NX9490 NX9461NX9463

NX9464

NX94121NX94120

NX94112

NX94106

Figure 4-6: Sample Location map for Naxos(adapted from Jansen, 1973; Buick, 1988).


50 µm

NX9451d.

234

grain 26

89646

731

c.

grain 150 µm

50 µm

244

SY9603

grain 11

b.a. SY9603

100 µm

e.

50 µmNX9461

g.

grain 2

3072

NX9490

100 µm

200 µm

f. NX9490

50 µm

NX94106h.

grain 19

307

Figure 4-7: CL images of zircon grains from Cycladic Series rocks: a) Regular, elongate, oscillatoryzoned zircons from sample SY9603; b) Age from core of oscillatory zones zircon from the Variorthogneiss (SY9603); c) Oscillatory zoned core of zircon in sample 89646 showing a Precambrian agefrom an irregular overgrowth; d) Regular oscillatory zoned grains with sharp terminations from sampleNX9451; e) Elongate, oscillatory zoned grains with low luminescent unzoned overgrowths (NX9461); f)Range in zircon morphologies from elongate, euhedral to rounded and squat from sample NX9490; g)Inherited Archaean core in a comlexly zoned zircon from NX9490; h) Inherited Permo-Carboniferous coreof oscillatory zoned grain (NX94106).

Chapter 4 101Analyses of thirty-one zircons from sample NX9451 yielded ages mostly in the

range ca . 250 to 220 Ma but with one analysis at ca . 138 Ma from an unzoned

overgrowth with low luminescence (Figure 4-8). One main zircon population can be

identified at 232 ± 2 Ma (n = 23) with an older population forming a slight shoulder on

the main peak at 243 ± 3 Ma (n = 7) from mixture modelling (Table 4-1). On the Tera-

Wasserburg Concordia diagram, the analyses form an array along a line extending back to

the common Pb composition (Figure 4-8) and exhibit no signs of Pb loss or U gain. The

zircons were generally unabraded, which indicates only a small amount of sedimentary

transport. This inference, combined with the Triassic ages, suggests that the quartzite is a

volcaniclastic sedimentary rocks derived from a proximal source.

0

2

4

6

8

10

12

14

16

0 50 100 150 200 250 300

NX9451n = 36

0

0.05

0.10

0.15

0.20

0.25

0.30

0 10 20 30 40 50 60

200 Ma300

No.

of

Ana

lyse

s

Age (Ma)

Common Pb


Analyses of fifteen zircons from sample NX9481, yield ages predominantly in the

range ca . 250 to 195 Ma, but with two inherited ages from grain 15 at ca . 315 Ma

(Appendix E), and three ages less than 175 Ma are also distinguishable with the youngest

age being ca . 120 Ma (Figure 4-9). Age groups for NX9481 can be distinguished at 210

± 2 Ma (n = 6), 228 ± 2 Ma (n = 9) and 241 ± 2 Ma (n = 6). The preponderance of

Triassic ages in both samples NX9481 and NX9451 suggests that they were derived from

a provenance dominated by Triassic rocks containing very little inheritance. This

conclusion is consistent with the sources for these samples being predominantly from

102 Triassic/JurassicTriassic igneous rocks. A minor contribution of material from a Permo-Carboniferous-

aged source, such as the Naxos Basement (Chapter 3) is also required.

0

0.05

0.1

0.15

207

Pb /

206

Pb

10 20 30 40 50 60

100 Ma200300

238 U / 206 Pb

0

1

2

3

4

5

6

100 150 200 250 300 350

NX9481n = 26

No.

of

Ana

lyse

s

Age (Ma)

Figure 4-9: Combined histogram with ~ 7 Ma bin widths and kerned probability density curve forzircons from sample NX9481. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

In addition to the quartzites, eight samples of calc-silicates and pelites of mixed

metamorphic grade from the Mesozoic Series rocks of Naxos were analysed. Analyses of

five zircons from NX9461, a quartz-calcite-phengite schist from the Diaspore Zone of

Naxos (Jansen, 1973; Jansen and Schuiling, 1976), yield an age range from ca . 620 to

110 Ma, along with two Cretaceous ages (Figure 4-10). The zircons giving the oldest

ages are a mixture of abraded, sedimentary grains with pitted surfaces and rounded

morphologies. In contrast, the Cretaceous ages are from unabraded, elongate, zircons

preserving sharp terminations (Figure 4-11). Both grain types displayed regular

oscillatory zoning with overgrowth and core structures evident (Figure 4-7). The

youngest protolith age is from the core of an oscillatory-zoned grain aged ca . 623 Ma.

However, the abraded appearance of the grains yielding Silurian ( ca . 430 Ma) and

Jurassic ( ca . 166 Ma) ages suggests that sedimentation was Jurassic. The Cretaceous

ages are all from overgrowths and are interpreted to be of metamorphic origin (Chapter

5). The small number of analyses precludes a rigorous assessment of the timing of

sedimentation, although it appears to be Mesozoic.

Chapter 4 103

0

0.1

0.2

0.3

0.4

0 20 40 60 80 100

100 Ma300 200

207

Pb /

206

Pb

238 U / 206 Pb

0

1.0

2.0

0 100 200 300 400 500 600 700

NX9461

No.

of

Ana

lyse

s

Age (Ma)

n = 5


200 µm

NX9451a. NX9461a. NX94106b.

200 µm

90346d.

500 µm

c. IO9615

500 µm

Figure 4-11: Transmitted light photomicrographs of zircons: a) Range of zircon morphologies fromsample NX9461 ranging from rounded abraded grains with surface pitting to clear, elongate euhedral grainswith only slightly rounded terminations; b) Heterogeneous population of zircons from pelite NX94106containing grains showing evidence of sedimentary transport as well as magmatic-appearing grains; c) andd) Show a range in zircon morphologies similar to those described for (b), in Ios samples IO9615 and90346 respectively.


NX9463 is a low M2 grade quartz-calcite-chlorite-sericite schist also sampled from

the Diaspore Zone, and contains small (< 200 µm) zircons with similar morphologies and

internal structures as those described for NX9461. The range in ages determined by

analyses of fourteen zircons is ca . 3170 - 75 Ma (Figure 4-12) which is very large. The

youngest ages were obtained from analyses of rims on oscillatory zoned grains with the

youngest definite protolith age is from a oscillatory zoned zircon core at ca . 304 Ma. The

four youngest ages at ca . 75 Ma, are possibly new zircon growth associated with

metamorphism prior to M1. With the exception of one grain, all Cretaceous ages are from

homogeneous unzoned overgrowths and it is unclear to what extent these data represent

protolith ages or metamorphic ages, so the most that can be suggested is that the sediment

is no older than Permian and most likely Mesozoic in age.

NX9463

0

0.05

0.10

0.15

0.20

0.25

0.30

0.35

0 20 40 60 80 100 120

75 Ma100200300

n = 18

207

Pb /

206

Pb

238 U / 206 Pb

0

1

2

3

4

5

0 500 1000 1500 2000 2500 3000

No.

of

Ana

lyse

s

Age (Ma)


Chapter 4 105Sample NX94112 is a calc-silicate from the chlorite-sericite zone (Jansen and

Schuiling, 1976) that has experienced slightly higher temperatures during M2 than the

previous samples described. Zircon grains varied in size but were generally < 100 µm in

length, with rounded terminations indicative of abrasion or dissolution and internal

structures revealing homogeneous regularly zoned grains, commonly with fine (< 30 µm

in width) overgrowths that were unzoned and poorly luminescent. Ages obtained from

seven zircons ranged from 55 to 685 Ma, with most ages being Cretaceous (Figure 4-13).

All analyses, except those yielding the oldest and youngest ages, were determined from

clear overgrowths on oscillatory-zoned grains. The 690 Ma age is from one of the

oscillatory zoned grains, which showed no overgrowth structures. This age is interpreted

as a protolith age, placing a maximum constraint on the timing of sedimentation. The 55

Ma age was from the edge of a weakly oscillatory-zoned grain which had possibly

undergone recrystallisation/replacement as internal zoning was much more evident

towards the core of the grain, a feature typical of zircons which have undergone such a

process (Pidgeon, 1992). The other Cretaceous ages are interpreted as being of

metamorphic origin and are discussed in more detail in Chapter 5.

0

0.1

0.2

0.3

0 20 40 60 80 100 120 140

50Ma75100500

207

Pb /

206

Pb

238 U / 206 Pb

0

1

2

3

0 100 200 300 400 500 600 700 800Age (Ma)

No.

of

Ana

lyse

s

NX94112n = 7


106 Triassic/JurassicSample NX9464 is a calc-silicate with dominantly quartz-biotite-calcite mineralogy

from a high M2 grade section of Naxos (the Kyanite Zone of Jansen, 1973; Jansen and

Schuiling, 1976). It contains transparent, commonly broken zircons with rounded edges

and pitted surfaces. Internally grains show regular oscillatory zoning, with some

complex core and overgrowth structures (Figure 4-7). This complexity is reflected in the

age range of ca . 2860 to 180 Ma obtained for the sample from 31 zircons (Figure 4-14).

As all zircons have an abraded appearance and no new metamophic zircon overgrowths

are visible, the youngest ages suggest the time of sedimentation is Jurassic or younger.

NX9464n = 37

0

0.05

0.10

0.15

0.20

0.25

0.30

0 10 20 30 40 50

200 Ma300400

207

Pb /

206

Pb

238 U / 206 Pb

0

1

2

3

4

5

6

7

8

0 500 1000 1500 2000 2500 3000

No.

of

Ana

lyse

s

Age (Ma)


Chapter 4 107Sample NX94120 is also a calc-silicate consisting of quartz-calcite-biotite but is of

slightly higher metamorphic grade than NX9464, being from the Sillimanite Zone (Figure

4-6). It hosts a variety of complexly-zoned zircons. These zircons contain resorbed,

often irregularly zoned cores, overgrown by complex rims that are commonly spongy in

appearance and inclusion-rich (Figure 4-7). Analyses of ten zircons gave a range of ages

from ca . 250 to 30 Ma. The core yielded the older ages while the younger ages were

restricted to rims grown during metamorphism (Figure 4-15). A distinct group of zircons

is apparent at 232 ± 3 Ma (n = 5) on the age probability density curve and from mixture

modelling (Table 4-1). The youngest protolith age was measured from an unzoned

xenocrystic core at ca . 162 Ma, which suggests that sedimentation occurred during the

Triassic or later.

0

1

2

3

0 50 100 150 200 250 300

NX94120n = 13

0

0.1

0.2

0.3

0.4

0 50 100 150 200 250 300

30 Ma405075100200

207

Pb /

206

Pb

238 U / 206 Pb

No.

of

Ana

lyse

s

Age (Ma)


108 Triassic/JurassicSample NX94121 is also from the Sillimanite Zone (Jansen, 1973; Jansen and

Schuiling, 1976) and is a quartz-rich calc-silicate with the assemblage quartz-

clinopyroxene-plagioclase-scapolite-calcite-titanite-garnet-epidote. As it was the first

sample from which new young metamorphic growth rims were identified, it was analysed

repeatedly and these rims were preferentially dated. Most of the zircons in the sample

showed evidence of new growth, although some grains had growth rims less than 10 µm

wide around generally uncomplicated oscillatory zoned cores. Such cores occasionally

preserved several periods of zircon growth. From one hundred and thirty-one analyses of

ninety-five zircons, a range in ages from ca . 1660 to 13 Ma has been established (Figure

4-16). A large number of Cretaceous metamorphic zircon growths are found in this

sample and are described in detail in Chapter 5, while the complexity of Tertiary-aged

metamorphic overgrowths is discussed in Chapter 6. The youngest protolith aged zircon

dated at ca . 213 Ma is an oscillatory-zoned zircon displaying no overgrowths. This age

therefore constrains the age of the sedimentary protolith of the metamorphic calc-silcate to

Jurassic or younger. Note that this sample also contains a notable number of inherited

Permo-Carboniferous Basement-aged zircons, suggesting that the source was comprised,

at least in part, of material having this age.

0

5

10

15

20

25

0 100 200 300 400 500 600 700 800

NX94121

0

0.1

0.2

0.3

0.4

0.5

0.6

0.7

15 Ma203040

0 100 200 300 400 500

207

Pb /

206

Pb

238 U / 206 Pb

n = 131

No.

of

Ana

lyse

s

Age (Ma)

Figure 4-16: Combined histogram with 10 Ma bin widths and kerned probability density curve forzircons from sample NX94121. Inset ia a Tera-Wasserburg Concordia diagram showing all analyses. Notethat one analysis at 1664 Ma is not included on the histogram or age probability curve.

Chapter 4 109Two samples of pelitic composition were analysed, one of low M2 grade and one of

high M2 grade from the Chlorite/Sericite and Migmatite zones respectively (Jansen, 1973;

Jansen and Schuiling, 1976). Sample NX9490 has a low grade quartz-albite-chlorite-

phengite mineralogy. It contains typically abraded, coloured grains that are well-rounded

with pitted surfaces. These detrital-appearing grains are interspersed with rare magmatic

grains which in contrast retain some angular terminations. Complex internal growth

structures are evidence of complicated growth histories and several rim structures are

visible enclosing older cores (Figure 4-7). The age range in forty-one zircons from

NX9490 was described in Chapter 1 and is illustrated in Figure 4-17. The youngest age

of ca . 45 Ma is from a low luminescent, low Th/U zircon growth rim, interpreted as a

metamorphic overgrowth. The next two youngest ages are Cretaceous and come from

zircon overgrowths, resulting from metamorphism. The next youngest age is ca . 373 Ma

which places a maximum age constraint on the timing of sedimentation of the protolith as

Devonian.

0

1

2

3

4

5

6

0 500 1000 1500 2000 2500 3000

NX9490

0

0.05

0.10

0.15

0.20

0.25

0.30

0 50 100 150 200

200 40 Ma5075100

207

Pb /

206

Pb

238 U / 206 Pb

n = 50

No.

of

Ana

lyse

s

Age (Ma)


110 Triassic/JurassicThe high-grade M2 pelite sample NX94106, was taken from a unit identified as a

pelitic “raft” in the leucogneiss core of Naxos (Buick, 1988) (Figure 4-6). This unit was

sampled to test whether it is related to the Mesozoic Series rocks which occur outside the

core, as has been previously suggested (Jansen and Schuiling, 1976). It contains the

assemblage K-feldspar-quartz-biotite-sillimanite and has a range of zircon morphologies

(Figure 4-11) with typically resorbed oscillatory zoned zircon cores mantled by

overgrowths of metamorphic zircon (Figure 4-7). Analysis of twenty-seven grains

produced ages ranging from ca . 1040 to 16 Ma (Figure 4-18). Cretaceous and younger

ages were derived from clear, unzoned, low luminescent zircon overgrowths, and

probably record the timing of a metamorphic overprint on the sediments. The youngest

protolith age in the sample occurs at ca . 293 Ma, which constrains the timing of

deposition of the sedimentary protolith to be Permian.

n = 31NX94106

0.04

0.08

0.12

0.16

0.20

0 100 200 300 400 500

15 Ma203050100

207

Pb /

206

Pb

238 U / 206 Pb

0

2

4

6

8

10

12

0 200 400 600 800 1000 1200Age (Ma)

No.

of

Ana

lyse

s


Chapter 4 111

4.5.3 Ios

Three samples from the Ios Upper

plate (Forster and Lister, 1996), which are

assumed to be Mesozoic Series rocks

(Dürr, 1986), were analysed. The

locations are indicated in Figure 4-19.

These samples are: 89639 (a sample from

Baldwin, 1996), 90346 (from the RSES

ANU rock collection) and IO9615. 89639

is a glaucophane schist affected by M2

metamorphism (Baldwin, 1996). Sample

89639 contains a mixture of honey-

coloured, rounded, commonly broken,

abraded, irregularly-zoned zircons as well

as clear, colourless, euhedral oscillatory

zoned grains preserving sharp

terminations. Many of these zircons have

spongy overgrowths, marked by a trail of inclusions, which are concentrated at grain

terminations and overgrow variably zoned cores. The measured age range from ca . 355

to 60 Ma for thirty-two zircons reflects the heterogeneity of the zircon populations (Figure

4-20).

Despite the metamorphic appearance of the zircon overgrowths, none of the grains

analysed had the very low Th/U ratios characteristic of some metamorphic zircons

(Williams and Claesson, 1987). This feature cannot be considered as solely diagnostic of

metamorphic zircon however, as evidenced by the range in Th/U ratios reported for

hydrothermally-precipitated zircon and recrystallised zircon (Claoue-Long et al., 1992;

Yeats et al., 1996). For this reason, the young Palaeocene and Cretaceous ages found

from overgrowths in these samples are interpreted as representing an episode of

metamorphic zircon growth. The youngest ages NOT from overgrowths occur at ca . 235

Ma (89639), suggesting that the maximum age of sedimentation of the schist protolith is

Jurassic. This is the only sample from the Ios Upper plate which contains inheritance the

age of the Ios Basement at ca . 300 Ma, suggesting a partly Permo-Carboniferous aged

source for some of the zircons. With the exception of the single age at ca . 60 Ma, all the

zircon overgrowths are much older than the inferred timing of M1 metamorphism in the

Cyclades of ca . 54-50 Ma (Baldwin, 1996). 40Ar-39Ar dating of white micas from sample

89639 yields an age gradient from 25.3 to 49.3 Ma, with a total fusion age of 42.2 ± 0.5

Ma which has been interpreted as partial outgassing/recrystallisation of M1 white mica

0 2 4 kmN

OrthogneissGarnet-Mica Schist

MarbleSchist

90346

89639IO9615

Figure 4-19: Ios sample location map (adaptedfrom van der Maar and Jansen, 1983).

112 Triassic/Jurassic(Baldwin and Lister, 1998). This younger resetting event is not recorded in the zircons

from this sample.

89639n = 43

0

0.05

0.10

0.15

0.20

0.25

0 20 40 60 80 100 120 140

500 200 100 75 50 Ma20

7 Pb

/ 20

6 Pb

238 U / 206 Pb

0

1

2

3

4

5

0 100 200 300 400 500Age (Ma)

No.

of

Ana

lyse

s


Sample 90346 is a quartz-phengite schist containing zircons with heterogeneous

morphologies ranging from elongate and euhedral to rounded and pitted (Figure 4-11).

All grains are consistently clear and colourless. As found in 89639, a range of irregular

to oscillatory zoned cores are overgrown by rims with new growth concentrated at the

terminations of grains. Ages from cores and rims (40 analyses of 38 grains) range from

ca . 1050 to 30 Ma (Figure 4-21). Four Eocene ages are found and these are interpreted as

possible metamorphic ages along with Palaeocene-aged overgrowths at ca . 60 Ma. The

age of the sedimentary protolith, as constrained by the youngest detrital zircon, occurs at

ca . 214 Ma (90346) suggesting that, like 89639, the sample is Jurassic or younger.

Chapter 4 113

0

2

4

6

8

10

12

0 200 400 600 800 1000 1200

90346

0

0.1

0.2

0.3

0.4

0.5

0 20 40 60 80 100 120 140500 200 100 75 50 Ma

n = 40

238 U / 206 Pb

207

Pb /

206

Pb

No.

of

Ana

lyse

s

Age (Ma)


IO9615

0

0.05

0.10

0.15

0.20

0.25

0.30

0 50 100 150

1000

500 100 75150 50 40

n = 24

238 U / 206 Pb

207

Pb /

206

Pb

0

2

4

6

8

10

0 100 200 300 400 500 600 700

200

No.

of

Ana

lyse

s

Age (Ma)

Figure 4-22: Combined histogram with 20 Ma bin widths and kerned probability density curvefor zircons from sample IO9615. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

114 Triassic/JurassicSample IO9615 is a garnet-glaucophane schist which, like the other two Ios samples,

contains a range of heterogeneous zircons with different shapes, sizes and internal

structures. Most grains appear to have undergone sedimentary transport, showing

surface pitting, although not all are rounded and some retain moderately elongate,

euhedral morphologies (Figure 4-11). Overgrowths in most zircons are commonly

concentrated at grain terminations. Ages obtained from twenty-one zircon grains range

from ca . 650 to 40 Ma (Figure 4-22). The youngest age not derived from a zircon

overgrowth is ca . 205 Ma, and was obtained from a clear oscillatory-zoned grain. This

result is in good agreement with the sedimentation ages for the other Ios Series rocks and

suggests sedimentation occurred in the Jurassic or younger. If the overgrowths are

actually metamorphic then they constrain the minimum timing of sedimentation to be

Cretaceous.

4.5.4 Folegandros

FL9602 is a pelite from the Mesozoic

Series rocks of Folegandros (location

Figure 4-23). It is a strongly foliated

quartz-chlorite schist with minor calcite,

and shows textural evidence of garnets

entirely pseudomorphed by quartz and

chlorite.

The sample contains a typically mixed

detrital population of zircons consisting of

rounded, honey-coloured, irregular and

oscillatory zoned grains, as well as more

angular, elongate to euhedral, colourless

and oscillatory zoned grains. Analyses of

nineteen zircons indicate a range of ages

from ca . 2880 to 90 Ma grains (Figure 4-24).

The youngest age in this population has a large uncertainty due to its relatively high

common Pb content (Appendix E). The next youngest age is ca . 140 Ma, and it is taken

to represent the protolith age for the Folegandros schist. A number of zircons contain

Triassic inheritance at 225 ± 3 (5) Ma (Table 4-1).

0 1 2 km

N

Marble

Schist

Alluvium

FL9602

Figure 4-23: Folegandros sample location map(adapted from Verginis, 1973).

Chapter 4 115

FL9602

0

0.05

0.10

0.15

0.20

0.25

0.30

0 10 20 30 40 50

200 Ma3005001000

3000n = 28

238 U / 206 Pb

207

Pb /

206

Pb

0

1

2

3

4

5

6

7

0 500 1000 1500 2000 2500 3000

No.

of

Ana

lyse

s

Age (Ma)

Figure 4-24: Combined histogram with 50 Ma bin widths and kerned probability density curve forzircons from sample FL9602. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

4.5.5 Sikinos

Sample SK9603 is a metabasic schist

from the Mesozoic Series rocks of Sikinos,

shown in Figure 4-25. Its zircons are

irregularly-shaped, pitted and contain

numerous inclusions. They commonly

display oscillatory zoning with rare inherited

cores (Figure 4-29). Twenty-one analyses

of seventeen zircon grains yield ages ranging

from ca . 230 to 45 Ma (Figure 4-26). The

youngest age is from a zircon with an

unusually high uranium content which grew

in a fracture in a pre-existing zircon grain. It

is possible that this grain could be new

metamorphic zircon growth. The next youngest zircon grain at ca . 60 Ma is from a zircon

with a mottled appearance, possibly due to new growth or recrystallisation, whereas the

other few ages younger than 180 Ma are from unzoned portions of otherwise oscillatory-

zoned grains. The first clearly magmatic-appearing, oscillatory-zoned zircon is 180 Ma in

age, placing a maximum constraint on the timing of sedimentation. This sample also

0 1 2 km

Orthogneiss

Marble

Garnet Mica Schist

Schist

N

SK9603

Figure 4-25: Sikinos sample location map (after Franz et al., 1993).

116 Triassic/Jurassiccontains a dominant population of Jurassic-Triassic inheritance at 201 ± 2 (4) and at 223

± 2 (5).

238 U / 206 Pb

0

1

2

3

4

5

0 50 100 150 200 250 300

SK9603

No.

of

Ana

lyse

s

Age (Ma)

n = 21

0

0.05

0.10

0.15

0.20

0 20 40 60 80 100 120 140

50 Ma75100200207

Pb /

206

Pb

Figure 4-26: Combined histogram with 10 Ma bin widths and kerned probability density curve forzircons from sample SK9603. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

4.5.6 Sifnos

A sample of a quartz-rich calc-silicate (SIF9345) was collected from the greenschist

unit of Sifnos, which comprises schists and marbles that are equivalent to the Mesozoic

Series rocks described on other islands (Schliestedt and Matthews, 1987). Sample

SIF9345 is a quartz-calcite-muscovite-albite schist containing zircons of variable size and

shape, that are generally colourless and occassionally inclusion-rich. Some zircons

appear detrital with obvious signs of abrasion such as rounded edges and pitted surfaces,

whereas other grains are angular with euhedral shapes. Most grains contain some sort of

core structure, generally with oscillatory zoning (Figure 4-29) and are commonly

overgrown by spongy rims mantled by inclusion-free outer rims. Forty-nine analyses of

forty-four zircons yielded a range of ages from ca . 2670 to 25 Ma (Figure 4-28).

Chapter 4 117

0 1 2 km

N

Marble

Schist

Alluvium

EBD

GSD

SIF9345

Figure 4-27: Sifnos sample location map (from Schliestedt and Okrusch, 1988).

SIF9345n = 49

0

0.2

0.4

0.6

0.8

0 50 100 150 200 250

207

Pb /

206

Pb

0

2

4

6

8

10

0 100 200 300 400 500 600 700

No.

of

Ana

lyse

s

Age (Ma)

238 U / 206 Pb

50 30 Ma

Figure 4-28: Combined histogram with 10 Ma bin widths and kerned probability density curve forzircons from sample SIF9345. Inset is a Tera-Wasserburg Concordia diagram showing all analyses. Notethat three analyses at 1693, 2022 and 2667Ma are not included on the histogram or age probability curve.


50 µm

NX9451a.

200 µm

SK9603a.

grain 2

grain 4

grain 6

grain 7

220

186 180220 185

201

50 µm

SIF9345b.

310

grain 32

Figure 4-29: Cl images of zircons from samples: a) SK9603 and b) SIF9345.

On the basis of overgrowth morphology all ages less than 200 Ma are considered to

be new metamorphic overgrowths. The youngest potential protolith age occurs at ca . 242

Ma derived from the core of a clear, oscillatory zoned grain. This age constrains the

timing of sedimentation to Triassic or younger. The presence of ca . 300 Ma detrital

zircons in the Sifnos calc-silicate suggests that some of the sediments were eroded from

Permo-Carboniferous basement.

Chapter 4 119

Table 4-1: Summary of U-Pb ages for the Vari Gneiss and Mesozoic Series samples

Sample No.spots

No.zircons


Age Range

SY9603 25 25 orthogneiss 240 ± 1 (25) 211-25689646 27 26 quartzite 288 ± 2 (7) 76 - 650

SY9630 23 21 quartzo-feldspathicschist

135 ± 2 (4)199 ± 1 (6)221 ± 1 (9)

107 - 271

NX9451 36 30 quartzite 232 ± 2 (23)243 ± 3 (7)

138-249

NX9481 26 15 quartzite 210 ± 2 (6)228 ± 2 (9)241 ± 2 (6)

118-316

NX9461 5 5 calc-silicate 112 - 623NX9463 18 14 calc-silicate 75.1 ± 0.8 (4) 74 - 3170NX94112 7 6 calc-silicate 55 - 685NX9464 37 31 calc-silicate 569 ± 4 (8)

647 ± 4 (8)180 - 2860

NX94120 13 10 calc-silicate 232 ± 3 (5) 29 - 251NX94121 131 95 calc-silicate 63.0 ± 0.9 (8)

96.1 ± 1.1 (8)120 ± 3 (6)142 ± 2 (10)168 ± 2 (5)323 ± 5 (5)

14 - 1665(note: Tertiaryages from this

sample are listedin Chapter 6)

NX9490 50 41 pelite 617 ± 6 (5)759 ± 12 (5)831 ± 8 (5)

2474 ± 15 (5)

45 - 3190

NX94106 31 27 pelite 18.3 ± 0.2 (7)20.6 ± 0.5 (4)

122 ± 1 (3)

16 - 1045

89639 43 32 glaucophane schist 88.8 ± 1.8 (4)191 ± 2 (5)232 ± 3 (4)275 ± 3 (4)287 ± 3 (4)

60 - 491

IO9615 24 21 garnet-glaucophaneschist

74.0 ± 2.0 (5) 40 - 645

44 40 38 quartz-phengiteschist

60.3 ± 0.7 (8)77.1 ± 0.7 (5)

127 ± 2 (4)152 ± 3 (6)

60 - 1050

FL9602 28 19 pelite 225 ± 3 (5) 91 - 2880SK9603 21 17 metabasic schist 182 ± 2 (5)

201 ± 2 (4)223 ± 2 (5)

47 - 228

SIF9345 49 40 calc-silicate 43.7 ± 0.9 (5)52.8 ± 0.8 (7)95.3 ± 1.8 (5)

35 - 2870


4.5.7 Combined Triassic-Jurassic U-Pb Zircon Ages

Figure 4-30 combines all the Triassic-Jurassic zircon ages from all samples

analysed from the Cyclades and discussed in this study. Figure 4-30 illustrates that the

main period of zircon formation occurred where the broad double peak is located between

250 - 220 Ma. This distribution indicates that 250-220 Ma was a period of major tectonic

activity in the Aegean, creating crustal material from which the Mesozoic Series rocks of

the Cyclades were sourced. There are three smaller peaks visible in the age probability

density diagram (Figure 4-30) at around ca . 220-200 Ma, 170 Ma and 150 Ma.

Interestingly, in at least two of the Triassic-aged samples analysed, the ages come from

very silica-rich rocks that contain distinctive zircons having only a single age peak of

approximately 220 Ma. These zircons show little evidence of transport which suggests

that they were derived from a proximal source and the sedimentary protoliths are assumed

to be volcaniclastic units derived from erosion of nearby volcanics.

0

5

10

15

20

25

140 160 180 200 220 240 260Age (Ma)

No.

of

Ana

lyse

s

Figure 4-30 : Combined histogram with 5 Ma bin widths and kerned probability density curve for allsamples discussed in this thesis, comprising 229 analyses to illustrate the major age populations.

Chapter 4 121

4.6 Discussion

4.6.1 Age of the Vari Gneiss, Syros

SHRIMP U-Pb dating of the Vari Gneiss sheds new light on the formation of this

metamorphosed igneous complex, previously thought to be Cretaceous in age. As

discussed in Section 4.2 and 4.3, the age of the Vari gneiss has been difficult to resolve.

However, SHRIMP U-Pb data from one sample of the Vari gneiss has identified an

essentially single population of zircons, free from the complications of inheritance, from

one intrusion comprising the Vari unit, indicating that it was emplaced at ca . 240 Ma. A

Triassic age for part of this deformed and metamorphosed igneous complex is an

unexpected result and indicates a period of magmatism in the Cycladic region during the

Triassic that was previously unrecognised. This has implications for the age of the

surrounding Syros schists which had been presumed to be significantly older than the

gneiss. The postulated existence of a contact aureole around the Vari gneiss (Ballhaus,

pers comm) places doubt on previous interpretations of a tectonic contact between the

gneiss and surrounding metasediments (Bonneau et al., 1980; Bonneau, 1984; Ridley,

1984). This would constrain the age of the Syros schists surrounding the Vari gneiss to

being older than the Vari gneiss, so Triassic, or most probably older. This is consistent

with age data for the Series rocks of the Cyclades, which range from Permian to Tertiary

in age (Dürr, 1986). Another implication of this interpretation is that if the Vari gneiss

protolith has intruded the Syros schists, which have demonstrably experienced Eocene M1

(Baldwin, 1996), then the Vari gneiss must have undergone the same metamorphic

history, making the preservation of Cretaceous-aged phengites (Maluski et al., 1987)

difficult to explain. Such Cretaceous ages are also found in rocks from the Upper Unit of

the Cyclades (which have not undergone M1) and these ages are thought to reflect the

timing of amphibolite-grade metamorphism and ophiolite emplacement (Reinecke et al.,

1982; Patzak et al., 1994). However, until the existence of a tectonic contact between the

Vari gneiss and the Syros schists is comprehensively proved or disproved, the full

significance of the age results will be difficult to assess.

4.6.2 Age of Series rocks

The provenance ages for metasedimentary Series rocks presented in this chapter

indicate that most were deposited in the Triassic-Jurassic or possibly later. This is

consistent with fossil evidence of Triassic deposition (Dürr et al., 1978) and also with the

age of leucogneiss protoliths from the Naxos core (Buick, 1988) that also yielded

Triassic-Jurassic provenance ages (Chapter 3). The samples described in this chapter

were interpreted as belonging to the Mesozoic Series rocks rather than the Basement that

is predominantly Permo-Carboniferous in age. The presumed Eocene timing of M1

122 Triassic/Jurassicmetamorphism (Altherr et al., 1979; Andriessen et al., 1979; Wijbrans and McDougall,

1988; Baldwin, 1996) places a minimum age constraint on the time of sedimentation,

while the recogonition of Cretaceous-aged metamorphic zircon overgrowths suggests that

the rocks have experienced a complicated pre-M1 history.

For the Series rocks to have experienced metamorphism in the Cretaceous, they

must have been sediments deposited prior to the Cretaceous. While it is possible that they

were then uplifted and eroded to form younger sedimentary sequences prior to Eocene M1

metamorphism, such an explanation allows little time for the operation of major tectonic

processes and is unsupported by the zircon morphology (Cretaceous-aged rims show little

evidence of subsequent sedimentary abrasion). For these reasons, it seems more likely

that the sediments of the Series rocks were deposited in the Triassic-Jurassic and

underwent pre-M1 metamorphism of unknown grade which may be correlated to the

timing of Eo-Alpine metamorphism in other areas of the Alpine mountain chain (as will be

discussed in Chapter 5).

Table 4-1 illustrates some of the main Triassic-Jurassic age populations present in

zircons from the metasedimentary Series rocks. While a few samples preserve large

populations of Pan-African and Neoproterozoic ages (with one sample preserving a

number of Paleoproterozoic ages), the younger ages reflect sources distinguished in

Chapters 3 and 4. Sample NX94121 and NX9481 preserve zircon age populations

consistent with derivation from a source similar to the 330-300 Ma basement

orthogneisses described in Chapter 3 while samples from Syros and Ios (89646 and

89639) preserve populations in the ca . 290 - 275 Ma age range that could also be related

to Basement ages. These ca . 290 - 275 Ma ages for Basement units have also been

suggested on the basis of zircon Pb evaporation dating from Naxos orthogneiss samples

(Reischmann, 1998). The presence of strong 300 Ma populations in zircons from some

Series rocks suggests that regardless of whether the Series sediments were directly

deposited on Basement, Basement-aged material is required to contribute material to the

formation of the Series sediments. This suggests that the Cycladic Basement, or material

of the same age, was emergent prior to the Triassic.

The Series rocks of the Cyclades described in this chapter were derived from a

variety of different-aged sources (or sources comprised of heterogeneous age

populations) except for NX9481 and NX9451 which come from a predominantly single-

age source. Series sediments were most likely deposited in an active tectonic environment

similar to a modern-day back-arc basin (as inferred from ophiolite geochemistry, see

Chpater 5), before they underwent Cretaceous and then Tertiary metamorphism associated

with the Alpine orogeny.

4.6.3 Correlation with Menderes Massif, Turkey

Chapter 4 123Direct correlations between the Series rocks of the Cyclades and those of the

Menderes Massif have recently been made on the basis of lithostratigraphy and the similar

metamorphic histories of the two areas (Candan et al., 1997). While the exact location of

samples is not specified, Triassic ages have been reported for the Menderes Massif

(Dannat and Reischmann, 1997) and are thought to record a significant magmatic event at

ca . 240-230 Ma. Like the Cyclades, the Menderes Massif is considered to be a

metamorphic core complex formed during Alpine orogenesis (Bozkurt and Park, 1994;

Hetzel, 1995). The presence of Triassic ages in the Menderes Massif suggests that

correlations with the Cyclades may be accurate for Series rocks, if not for the Basement

units that they overlie (as discussed in Chapter 3). One possible difference between the

Cyclades and the Menderes Massif is that, except for the Cycladic island of Ios, where the

Series clearly form an upper plate in tectonic contact with a lower plate of Carboniferous-

aged Basement, the metamorphic core complexes of the Cyclades are generally identified

by detachments between Upper Unit and Series rocks (Lister et al., 1984) whereas the

contact between Basment and Series rocks is ambiguous (as discussed in Section 4.2).

This contrasts with the Menderes Massif, where the core complexes comprise

Precambrian basement exhumed along low angle detachments to juxtapose them against

Mesozoic Series rocks. It is conceivable that unexposed Precambrian basement

structurally underlies “Basement” of Carboniferous age exposed in the Cyclades and that

evidence of this latter unit has been eroded away in the Menderes Massif. If Permo-

Carboniferous material was eroded from the Menderes Massif after exhumation of the

Basement units then this could be tested by investigating the provenance ages of Pliocene

to Recent sediments formed during the exhumation process (these have a well-constrained

Mio-Pliocene age in the Cyclades (Roesler, 1978)) to see if a preponderance of

Carboniferous ages is obtained. If Permo-Carboniferous Basement existed in the

Menderes Massif but was eroded prior to exhumation then one might expect a

preponderance of Permo-Carboniferous ages in the Series rocks of the Menderes Massif

(Candan et al., 1997). If neither the Series rocks nor the Pliocene sediments of the

Menderes Massif contain zircons of this age, then some other explanation must be sought

to explain the apparent discontinuity between the Basement of the Cyclades and the

Menderes Massif, which most plate reconstructions suggest should be closely spatially

related (Robertson and Dixon, 1984).

4.6.4 Correlation with Pelagonian Zone, Internal Hellenides, Greece

Permo-Triassic silicic metavolcanics have been reported from the Pelagonian zone

(Mountrakis et al., 1987) as well as from the transitional area between the external and

internal Hellenides (Magganas et al., 1997). An early map displaying the widespread

distribution of Triassic volcano-sedimentary sequences in the Hellenides was presented

by (Celet et al., 1977). The development of passive margins accompanying Triassic

124 Triassic/Jurassicrifting and volcanism is clearly evident in the sub-Pelagonian zone where the

disintegration of a continental platform and onset of deep water sedimentation was

coupled with alkalic volcanism (Smith and Spray, 1984). The existence of Triassic aged

volcanic sources in close association with sedimentary sequences is a similar environment

to that proposed here for Naxos samples NX9481 and NX9451. If more of these

Triassic volcaniclastic sequences could be identified in the Cyclades this would be strong

evidence for a correlation between the Pelagonian zone and the Cyclades, but until further

evidence is available such a conclusion is speculative.

4.6.5 Correlation with External Hellenides

The Vari unit has been interpreted as a probable equivalent of the Asteroussia

Nappe of Crete (Seidel et al., 1976; Bonneau, 1984), which consists of an ophiolitic

melange of low-P high-T Cretaceous metamorphics with Jurassic-aged ophiolites (Seidel

et al., 1981). New SHRIMP U-Pb age constraints given in this chapter suggest that this

correlation is invalid. An age of ca . 240 Ma for the Vari gneiss suggests it is more likely

correlated with the Permo-Triassic Phyllite-Quartzite (PQ) unit of Crete (Seidel et al.,

1982) which consists of continental to shallow marine siliciclastics, including quartzites

as well as pyroclastic, volcanic and intrusive rocks (Bonneau, 1984; Hall et al., 1984).

The PQ unit has been likened to the Tyros unit of the Peloponessus, external Hellenides

(Bonneau, 1984) from which Triassic pyroclastics, basalts and andesites and dacites are

reported (Pe-Piper, 1982). Triassic magmatism in the PQ unit and in the Peloponessus is

contemporaneous with the formation of the granite protolith to the Vari gneiss of Syros

and also with the presumed volcanic protolith to samples NX9451 and NX9481

(described in Section 4.5). If these external Hellenide units are correlative with sections

of the Mesozoic Series rocks of the Cyclades then the volcanic and intrusive rock

associations found within them may once have existed in the Cyclades. This would

explain the abundance of Early Triassic ages found in the metasedimentary samples

analysed in this Chapter.

4.6.6 Tectonic Implications of SHRIMP ages

The new SHRIMP U-Pb zircon ages reported in this chapter provide evidence of

Triassic-aged sedimentation and volcanism in the Cyclades followed by Triassic-Jurassic

sedimentation. The age of the magmatic complex that forms the Vari gneiss at ca . 240 Ma

distinguishes a period of previously unrecognised magmatic activity in the Cyclades.

While volcanic units have been identified interlayered with the Cycladic Series rocks, no

age constraints on these units have previously been reported. The dominant Triassic age

of zircons in samples NX9451 and NX9481 that presumably represent immature volcanic

detritus places an indirect age constraint on volcanism in the Cycladic region that was

previously unrecognised. There is widespread evidence of sedimentation and volcanism

Chapter 4 125reported elsewhere in the Mediterranean region at this time (e.g., Greece, Turkey,

Albania, Yugoslavia, see Section 4.1). This tectonic activity is generally ascribed to

continental rifting and/or back-arc extension (e.g., Magganas et al., 1997). This rifting is

interpreted to occur in response to the opening of the Tethys ocean (Robertson and

Dixon, 1984; Mountrakis, 1986; Mountrakis et al., 1987), which is indicated by the

formation of the Pindos ocean and spreading centres south of the South Aegean block and

Menderes Massif (Figure 4-31).

Adriatic-Apulian Promontory

Moesia

Rhodope/Serbo-Macedonia

Kirsehir

Menderes/Tauride

E. Tauride

PuturgeBitlis

South Aegean

EURASIA

AFRICA

Alanya

Pelagonian

TRIASSIC

Robertson and Dixon (1984)

Greater CaucasusSubduction Zone

Lesser Caucasus

ArabianPromontory

PALAEOTETHYS

Tornquist-Teisseyre Lineament

~ 250 Ma

Figure 4-31: Triassic reconstruction showing the location of the South Aegean block (which includesthe Cyclades) as part of North Africa (then Gondwana), and also the proximity of the southern Eurasianmargin (Robertson and Dixon, 1984).

126 Triassic/JurassicThis rifting lead to the detachment of the various continental blocks delineated in

Figure 4-31, from the northern margin of Africa (then Gondwana), and their transport

towards the southern margin of Eurasia (Figure 4-32).

Vardar/PontideOphiolites

Eurasian Margin Collage

PindosOcean

Kirsehir

Sakarya

Menderes/Tauride

E. Tauride

Puturge

Bitlis

South Aegean

EURASIA

AFRICA

Alanya Robertson and Dixon (1984)

Pelagonian

MIDDLE JURASSIC ~180 Ma

Spreading centres

Figure 4-32 : Jurassic plate reconstruction showing the location of the Cyclades as part of the SouthAegean block, and the development of active spreading ridges forming deep rifts floored by stretchedcontinental crust, consistent with this being a period of active volcanism (reconstruction from Robertsonand Dixon, 1984).

These blocks eventually collided with Eurasia during the Cretaceous-Tertiary Alpine

orogeny. Evidence from the dating and correlation of magnetic anomalies on the Atlantic

ocean floor suggests that the cause of the initial rifting along the northern margin of Africa

Chapter 4 127was related to successive phases of relative shear and compression during the Mesozoic

the between Gondwanan and Eurasian land masses. The compression and closure of

ocean basins in the wake of these continental blocks coincides with the Late-Triassic-

Early Jurassic opening of the North Atlantic (Livermore et al., 1986) and the widespread

Triassic volcanism developed in the Alpine region is thought to be evidence of this

process (Hynes, 1974).

4.7 Synthesis

SHRIMP U-Pb dating of zircon from the magmatic complex forming the “Vari”

orthogneiss of Syros reveals a ca . 240 Ma (Triassic) magmatic crystallisation age for this

formerly problematic sample. Zircon ages from the Series rocks sampled from the islands

of Syros, Naxos, Ios, Folegandros, Sikinos and Sifnos reveal that the provenance ages

of these samples are dominantly Triassic-Jurassic, with a Permo-Carboniferous

“Basement”-aged source also identifiable. The fact that most of the zircons from these

samples also show Cretaceous-aged metamorphic overgrowths, which themselves show

little evidence of abrasion, suggests that sedimentation must have occurred prior to this

time. This constrains the timing of sedimentation to Late Jurassic/Early Cretaceous for

most samples and is consistent with the recognition of these units as broadly “Mesozoic”

(Dürr, 1986). The meaning and significance of these Cretaceous-aged zircon rims are

discussed in Chapter 5. The age for the Vari gneiss was unexpected. The Vari gneiss

protolith is much younger than the orthogneisses of the Cycladic basement (330-300 Ma

from Chapter 3), and falls within the age range of both the Cycladic Series and Upper

Unit. Whether the Vari gneiss forms part of one of these units is currently obscured by

ambiguous field relationships. If it intruded the surrounding Syros schists and

experienced M1 metamorphism, it must be assigned to the Cycladic Series. Regardless of

which unit the Vari gneiss belongs, the existence of granitic material of Triassic age in the

Cyclades and the predominance of Triassic-Jurassic-aged magmatic zircons in the Series

sediments is evidence of extensive magmatic activity in the Aegean region at this time.

This is consistent with the timing of rift development and concentrated magmatic activity

in surrounding regions such as the Menderes Massif of Turkey, the Pelagonian zone of

mainland Greece, the External Hellenides of Greece and the Dinarides of Yugoslavia.

Such widespread magmatic activity has been related to rifting and volcanism associated

with the early stages of the opening of the central Atlantic further west (Robertson and

Dixon, 1984).

Chapter 5 129

5. CRETACEOUS GEOLOGICAL EVOLUTION OF THE CYCLADES

5.1 Introduction

In the Cretaceous, an East-West Tethyan oceanic realm separated the continental

land masses into a northern and southern group, with Tethys merging at both ends with

the giant ocean Panthalassa - the proto-Pacific ocean (Figure 5-1). At this time,

Gondwana had begun to break-up with India, and Australia-Antarctica had separated from

South America-Africa. By the mid-Cretaceous a complete break was established between

North America-Eurasia and South America-Africa associated with a major change in the

relative motion of Eurasia relative to Africa (Robertson and Dixon, 1984). This change in

motion, from slow sinistral shear to rapid convergence of the African and Eurasian plates,

coincided with the opening of the North Atlantic at around 108 Ma.

Equator EastTethys

WestTethys

NorthAmerica Eurasia

SouthAmerica

Africa

India

AntarcticaAustralia

Pantha

lassa

Pantha

lassa

EARLY CRETACEOUS ~ 130 Ma

Figure 5-1: Early Cretaceous reconstruction of the continents showing the break-up of Pangea andseparation of Laurasia from Gondwana. The position of the Cyclades is not known precisely but isprobably just north of the northern margin of Africa and Arabia. (Figure adapted from Ricou, 1994).

Fragments of both continental and oceanic crust in the Aegean region have behaved

independently of the African and Eurasian plates since the Jurassic (Burchfiel, 1980).

The effects of Gondwana break-up on the Cyclades were the continued rifting of

microcontinental blocks from the northern margin of Gondwana and collision of these

blocks with Eurasia, resulting in closure of oceanic basins in their path and formation of

new oceanic basins in their wake (Robertson and Dixon, 1984). During the Cretaceous,

the spreading ridges between continental blocks and the northern margins of Africa and

130 CretaceousArabia were all subject to compression and collapsed, leading to the emplacement of

ophiolites on continental crust (Robertson and Dixon, 1984).

Remnants of oceanic crust, ophiolites, can provide important information about the

tectonic history of orogenic belts. Spectacular outcrops of eclogite-blueschist facies

metaophiolitic rocks are found on the Cycladic island of Syros (Ridley, 1984b) and

Sifnos (Okrusch et al., 1978; Schliestedt, 1986), with Syros being the type locality for

the high-P mineral glaucophane (Hausmann, 1845). Dating these metamorphosed

ophiolites places important constraints on the time of formation of the oceanic crust

precursor, but the crystallisation age of this material is often difficult to measure due to the

environment in which it forms and its mode of emplacement (Bill et al., 1997). The task

is made more complicated when the ophiolite has been subjected to high-P

metamorphism. Such is the case for the ophiolite units of the Cyclades, but dating of

these units has the potential not only to constrain the time at which oceanic crust

developed in the area, but also to constrain the maximum age of high-P metamorphism in

the Cyclades. The age of the ophiolites, and the timing of metamorphism in the

associated Series rocks of the Cyclades, are critical in the reconstruction of the complex

tectonic evolution of this region. This chapter examines the Cretaceous age of an

eclogitised ophiolite from the island of Syros, as well as discussing the Cretaceous

metamorphic ages found in the Cycladic Series rocks described in Chapter 4.


The ophiolites of Syros are thought to be contiguous with a belt of ophiolitic rocks

which can be traced across the Hellenides from southern Yugoslavia through mainland

Greece (Roddick et al., 1979) via Mt Olympos (Dercyke and Godfriaux, 1977) and

through the Attic-Cycladic into Turkey (Okay, 1984) (Figure 5-2). Two types of

ophiolite can be distinguished by their present-day locations, either in “Greek” areas

(Greece, Albania and former Yugoslavia), or “Turkish” areas (Turkey, Cyprus and

northern Syria), as defined by Robertson et al. (1996). They are referred to here as

“Greek” ophiolites and “Turkish” ophiolites, respectively. In broad terms the ophiolites

can be distinguished not only by their geographic position, but also by their age of

crystallisation and obduction onto continental crust. “Greek” ophiolites are typically

Triassic-Jurassic in age (Hynes, 1974; Roddick et al., 1979; Spray and Roddick, 1980;

Spray et al., 1984), the same age as most of the ophiolites in the western portion of the

Alpine chain (e.g., Bill et al., 1997), while “Turkish” ophiolites are typically Cretaceous

in age (Delaloye, 1977; Thuizat et al., 1981; Whitechurch et al., 1984). Ages from the

metamorphic “soles” of ophiolites typically post-date ophiolite formation by ~ 10-20 Ma

(e.g., Roddick et al., 1979; Ghent and Stout, 1981) and are thought to record the time of

ophiolite obduction. Variations in the ages of the metamorphic soles associated with

Greek and Turkish ophiolites can be related to different tectonic conditions. Late-

Chapter 5 131Triassic-Early Jurassic compression, the closure of “Greek” ocean basins and the

emplacement of ophiolites on continental crust is thought to be related to the opening of

the North Atlantic (Livermore et al., 1986), while “Turkish” oceanic basins which closed

in the Early Cretaceous are related to opening of South Atlantic (Robertson et al., 1996).

Crete

Black Sea

Eastern Mediterranean Cyprus

RHODOPE

MOESIANPLATFORM

CARPATHIANS

MENDERES

PANNONIANBASIN

PONTIDES

TAURIDES

Troodos

Levant

Crimea

0 100km

N

20º 30º

34º

40º

CYCLADES

VA

RDA

R

PELAG

ON

IAN

HELLENIDES

DINARIDES

TURKEY

GREECE

Figure 5-2: Distribution of ophiolites from mainland Greece through to Turkey (adapted fromRobertson et al., 1996).

Figure 5-2 highlights the distribution of ophiolites in the eastern Mediterranean

region, most of which represent relict oceanic crust preserving evidence of the opening

and closure of ancient oceanic basins. These outcrops have been crucial in generating

plate reconstructions. A progression in the timing of ophiolite sole metamorphism from

west to east is apparent, with eastward younging thought to record an eastward migration

of oceanic slicing related to closure of the Tethys ocean (Whitechurch et al., 1984).


There are few constraints on the age of the high-P ophiolites in the Series rocks of

the Cyclades, except for the fact that they must be older than the high-P metamorphism

(which is currently believed to be approx. 50-40 Ma) from radiometric dating (Altherr et

al., 1979; Andriessen et al., 1979; Henjes-Kunst and Kreuzer, 1982). Ophiolites of

Cretaceous age from sediments and ophiolitic melanges are located in the Cyclades at the

Upper Unit on Anafi (Reinecke et al., 1982), Tinos (Patzak et al., 1994), Nikouria and

132 CretaceousAmorgos (Dürr et al., 1978a), as well as south of the Cyclades in Crete (Seidel et al.,

1981), and north and west in Evvia and Argolis (Clift and Robertson, 1989; Robertson,

1991). These ophiolites have not experienced Alpine high-P metamorphism and have

never been correlated with the ophiolites of the Cycladic Series. Some of the ages are

reported as recording the timing of metamorphism during ophiolite emplacement, while

others, e.g., Anafi and Tinos, were considered to represent both the timing of

metamorphism and ophiolite formation. The age of the high-P ophiolites preserved in the

Cyclades remains poorly constrained (Robertson and Dixon, 1984). For this reason, a

sample of an eclogite boudin, representing part of metaophiolite sequence, was collected

from northern Syros for dating purposes.

Dating of high-pressure metagabbros and metasediments from Syros using K-Ar

systematics has been conducted by several workers. Blake et al. (1981) mention

40Ar/39Ar ages ranging from 80-40 Ma based on phengite, paragonite and glaucophane

from unspecified Syros samples. Maluski et al. (1987) report 40Ar/39Ar plateau ages of

varying reliability between 53-37 Ma derived from seven phengite concentrates. Two of

the samples investigated in this study were previously dated by Baldwin (1996) who

reports 40Ar/39Ar total fusion ages from white micas of 49.2 ± 0.2 Ma for a retrograde

eclogite (our sample 89642), 39.6 ± 0.1 Ma for a quartzite (our sample 89646), and 39.6

± 0.1 and 43.05 ± 0.12 from two retrograde blueschist samples. Baldwin (1996)

suggests that apparent ages in the range 54-50 Ma from blueschist and eclogite facies

rocks from both Syros and Ios represent the timing of subduction zone metamorphism

and that younger ages from 49-25 Ma are the result of partial retrogression of high-P

assemblages during greenschist facies metamorphism. These ages for the timing of high-

P metamorphism are slightly older than previous estimates but are consistent with ages

from the Olympos region in the Pelagonian zone (Schermer et al., 1990). They tie in well

with general estimates of the timing of high pressure metamorphism in the Cyclades

deduced mainly from argon data (Altherr et al., 1979), although as noted by Ridley

(1984b) they represent minimum age constraints and the timing of high pressure

metamorphism could be considerably older.

This chapter will first reveal the Cretaceous age of the Syros ophiolite and then

discuss the Cretaceous age of metamorphism preserved by the zircons in the majority of

Series rocks described in Chapter 4.

Chapter 5 1335.4 Dating of the Syros Ophiolite

Zircons from the Syros ophiolite were

extracted from a sample in the ANU rock

collection, sample 89642 (collected by

Dr. S. Baldwin). The sample site in

northern Syros, shown in Figure 5-3

was later visited. Sample 89642 is a

retrograde eclogite with well-preserved

igneous textures, presumed to have

formed from an ophiolite protolith

(Baldwin, 1996). Zircons from this

sample are generally between 100-200

µm in length, with aspect ratios of

approximately 3:1, and display regular

oscillatory zoning with rare core

structures (Figure 5-4).

Ages of ca . 104 to 68 Ma were

derived from twenty-three analyses on

thirteen zircons (Figure 5-5). With the

exception of one outlier at 104 Ma, the

analyses defined age populations of 74.8 ± 1.2 Ma (n = 17) and 88.8 ± 4.1 Ma (n = 5),

when deconvoluted using mixture modelling (Appendix D).

200 µm

89642a.

50 µm

89642b.

87 Ma

grain 9

77 Ma

Figure 5-4: a) Grey-scale transmitted light photomicrograph of zircons from the Syros ophioliteillustrating their general euhedral magmatic morphologies: b) CL image of zircon grain 9 from the Syrosophiolite showing a core dated at ca . 87 Ma, surrounded by an oscillatory-zoned rim dated at ca . 77 Ma.

The large age population at ca . 75 Ma is taken to represent the time of emplacement

of the ophiolite magma, while the scattered older ages are possibly caused by inherited

radiogenic Pb. Both younger and older age populations come from zircon grains which

are morphologically and chemically very similar; the older ages do not come from

distinctive inherited cores. There is no correlation between age and uranium content so it

0 1 2 km

N

Metabasite

Orthogneiss

Marble

Schist

89642

Figure 5-3: Location of Syros Samples,map (adapted from Hecht, 1984).

134 Cretaceousis considered unlikely that the large younger age population from the sample is produced

by Pb loss. Due to the dynamic nature of the ocean floor environment, it is possible that

the ca . 89 Ma ages were incorporated from older pre-existing oceanic crust into the

ophiolite magma, in which case this small age group could be considered as inherited.

0

5

10

15

0 50 100 150 200 250

0

0.1

0.2

0.3

0.4

0.5

0.6

0 20 40 60 80 100

75 Ma100

Common Pb

207

Pb /

206

Pb

238 U / 206 Pb

89642n = 23

Age (Ma)

No.

of

Ana

lyse

s

Figure 5-5: Combined histogram with 10 Ma bin widths and kerned probability density curve forzircons from sample 89642. Inset is a Tera-Wasserburg concordia diagram showing all analyses.

As zircon is not a common accessory phase in mafic magmas due to the high

anticipated solubility of zirconium oxide in low silica melts (Watson and Harrison, 1983),

the interpretation of the Syros ophiolite zircon ages as being representative of the timing

of magmatic crystallisation could be questioned. Black et al. (1991) described the

dangers inherent in interpreting zircons from mafic dykes in the Vestfold Hills of

Antarctica as dating the timing of dyke emplacement. These workers recognised two

ways in which erroneous ages could be produced:

1) incorporation of xenocrystic zircons which would yield anomalously old ages; and

2) metasomatic alteration resulting in the conversion of baddeleyite (ZrO2) to form zircon,

via the introduction of silica-rich fluid or the breakdown of silica-bearing minerals such as

pyroxene, which would yield anomalously young ages.

There are certain morphological indications which can help to differentiate whether

the zircons from the Syros ophiolite are the products of either magmatic crystallisation,

metasomatic alteration, or represent xenocrysts derived from crustal material. In general,

magmatic zircons are elongate with simple prismatic and pyramidal faces, while those

produced by metamorphism or metasomatism are more complexly faceted (Davis et al.,

Chapter 5 1351968). In contrast to zircons from the Syros ophiolite, metasomatic zircons from the

Vestfold Hills dykes are all very small, i.e. < 30 microns across (Black et al., 1991).

Xenocrystic grains can have a variety of morphologies and may be expected to show

some evidence of dissolution in a mafic magma due to the zircon undersaturation and

relatively high temperatures of such melts (Watson and Harrison, 1983). The zircons

separated from the Syros ophiolite are generally 100-200 µm in length, elongate, with

typical aspect ratios of 3:1, well-formed pyramidal terminations and relatively few facets.

Furthermore, most of the zircons contain inclusions of needle-like apatite crystals

consistent with crystallisation from a melt. Though some grains are broken, there is no

evidence of zircon dissolution in the form of rough surfaces or embayments in the grains,

as might be expected to occur after their incorporation into a mafic magma. In CL

images, the Syros ophiolite zircons display regular, oscillatory zoning with no inherited

cores or truncated growth zones evident. Although xenocrystic zircons might also display

these internal features, metamorphic and hydrothermal zircons, in particular, do not

usually display regular oscillatory growth zoning (eg.Claoue-Long et al., 1990).

The morphology and internal textures of zircons from the Syros ophiolite (Figure 5-

4) are typical of igneous zircons, and the lack of zircon inheritance is consistent with the

zircons being produced during one magmatic episode. If they are xenocrysts from a

sedimentary rock, they show no evidence of having either undergone a sedimentary cycle,

or of dissolution that would suggest a xenocrystic origin. The observation that, in thin

section, zircons are not associated with other Zr-bearing phases, such as baddeleyite

which may react to form zircon, suggests that zircon was not produced by metamorphic

processes after emplacement, unless Zr-bearing fluids were introduced into the rock.

However, a hydrothermal origin is not supported by the zircon morphology. The

consistency in U-Pb ages and the lack of evidence for later overgrowths suggests that the

zircons were not influenced by subsequent hydrothermal or metamorphic processes.

Further evidence supporting a magmatic origin for the Syros ophiolite zircons is presented

in the next section describing the zircon geochemistry.

5.4.1 Zircon Geochemistry

To help pinpoint the origin of zircons from the Syros ophiolite, the within-grain

trace element chemistry of eight zircons was determined by Paul Hoskin, Research

School of Earth Sciences, using SHRIMP in the energy-filtering configuration described

by Hoskin (in press). The results are listed in Table 5.1. The restricted range in Th/U

ratios found in the Syros ophiolite zircons (0.38 - 0.73) is consistent with average values

for granite zircons (0.15 - 1.20) (Ahrens et al., 1967) and also for mafic magmas (0.28 -

1.17) (Heaman, 1990). As the Th/U ratios are greater than 0.1, a magmatic rather than

metamorphic origin is favoured for these grains. As will be described in Chapter 6, the

metamorphic zircons found in this study typically have Th/U ratios much less than 0.1,

136 Cretaceousalthough higher Th/U ratios are found in Cretaceous metamorphic overgrowths described

later in this chapter. A range of Th/U ratios has been reported for zircons produced by

hydrothermal processes (Claoue-Long et al., 1990; Black et al., 1991), so Th/U ratios

alone cannot be used to rule out the possibility that the Syros ophiolite zircons were

produced by metasomatism.

Table 5-1: Zircon trace element chemistry in ppm, determined by SHRIMP.

Elements 1.1 2.1 3.1 4.1 5.1 6.1 7.1 8.1

P 1394 628 797 646 657 305 890 1021Y 7185 4518 6290 5752 4563 1175 2295 3714La 0.029 0.014 0.043 0.116 0.145 0.058 0.007 0.014Ce 53 9 12 11 11 4 22 30Pr 0.29 0.27 0.41 0.51 0.45 0.05 0.06 0.14Nd 5 5 8 9 6 1 1 2Sm 12 10 19 19 12 2 4 6Eu 2.5 2.5 4.5 4 2.6 0.4 0.9 1.3Gd 67 52 87 86 60 9 20 37Tb 40 28 45 43 32 6 12 20Dy 561 360 552 514 394 84 168 276Ho 236 151 213 201 159 37 72 117Er 1178 737 974 912 725 187 374 587Tm 254 154 192 179 152 41 83 125Yb 2099 1266 1569 1459 1218 384 721 1068Lu 414 250 298 281 239 73 146 244Hf 13447 10122 9756 10473 10819 10843 13048 12248Th 56 74 73 68 52 12 15 32U 133 101 112 98 93 32 23 61tot. REE 12107 7542 10264 9470 7574 2003 3920 6227

Th/U 0.42 0.73 0.65 0.69 0.56 0.38 0.68 0.53Hf wt% 1.34 1.01 0.98 1.05 1.08 1.08 1.30 1.22Ce/Ce* 142 34 23 11 10 17 253 164Eu/Eu* 0.27 0.32 0.34 0.30 0.29 0.30 0.32 0.26

A graphical representation of the chondrite-normalised REE values is given in

Figure 5-6. The results indicate that the zircons are strongly enriched in REE, with total

concentrations ranging from approximately 2000 to 12 000 ppm. A strong degree of

enrichment in the HREE is evident in Figure 5-6, which shows the zircons contain up to

10 000 times the value for chondrite (McDonough and Sun, 1995). This enrichment of

HREE relative to the LREE, and the presence of Ce and Eu anomalies is thought to be

typical of igneous zircon (Nagasawa, 1970) and, in particular, for zircon crystallised from

mafic melts (Heaman, 1990). Recent experimental work by Black and Hoskin (in prep.)

shows that, in contrast to igneous zircon, metamorphic zircon displays variable degrees of

enrichment in the LREE. This LREE enrichment is not evident in the REE patterns of the

Syros ophiolite zircons. Trace element variation diagrams illustrating the chemistry of the

Chapter 5 137Syros zircons are also displayed in

0.4 0.6 0.8 1.0 1.2 1.4 1.60

100

200

300

400

Lu

Hf wt%

500

Lu

/ Sm

Hf wt%0.4 0.6 0.8 1.0 1.2 1.4 1.60

10

20

30

40

50

0

0.5

1.0

1.5

2.0

2.5

3.0

0.4 0.6 0.8 1.0 1.2 1.4 1.6Hf wt%

Th

/ U

Kimberlites

Carbonatite and Nepheline Syenite

Mafic and Ultramafic samples

Felsic samples

Basalts

Syros Ophiolite Zircons

Figure 5-7.

0.01

0.1

1

10

100

1000

10000

100000

La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Syros Ophiolite Zircon

Cho

ndri

te N

orm

alis

ed

Rare Earth Elements

Figure 5-6: Chondrite normalised rare earth element plot for eight zircons from the Syros ophiolite.

138 Cretaceous

A chemical classification scheme has been developed by Heaman et al., (1990) to

describe the characteristic zircon chemistry found in different igneous rocks. Zircons

from the Syros ophiolite overlap and extend the field characteristic of zircons from

pyroxenites, gabbros and norites to higher Hf wt% levels (Figure 5-7). It should be

noted that the basalt field delineated by Heaman et al. (1990) is only for mafic volcanic

samples as opposed to mafic intrusives, and so it is widely separated from the

mafic/ultramafic fields shown in (Figure 5-7). The notably higher Hf wt% found in

Syros ophiolite zircons is not unusual. Due to its geochemical similarity to zirconium,

Hf4+ is the main substituent into the Zr4+ site in the zircon lattice and zircons typically

contain approximately 1 wt% Hf (Kinny, 1991), though igneous zircons with up to 4.8

wt% Hf have been reported (Rubin et al., 1989).

Chapter 5 139

0.4 0.6 0.8 1.0 1.2 1.4 1.60

100

200

300

400

Lu

Hf wt%

500

Lu

/ Sm

Hf wt%0.4 0.6 0.8 1.0 1.2 1.4 1.60

10

20

30

40

50

0

0.5

1.0

1.5

2.0

2.5

3.0

0.4 0.6 0.8 1.0 1.2 1.4 1.6Hf wt%

Th

/ U

Kimberlites

Carbonatite and Nepheline Syenite

Mafic and Ultramafic samples

Felsic samples

Basalts

Syros Ophiolite Zircons

Figure 5-7: Trace element variation diagrams for Syros ophiolite zircons overlain on fields recognisedfor igneous zircon suites (according to Heaman et al. 1990).

The trace element chemistry of the zircons is consistent with the whole rock

chemistry for the ophiolite sequences reported by Seck et al. (1996). The Syros ophiolite

zircons overlap with the Th/U values for felsic samples, but this is not uncommon in

zircons from mafic magmas (Heaman, 1990). To reiterate the major points raised in this

discussion, the general morphology and trace element composition of zircons from the

Syros ophiolite support their origin as a primary crystallisation phase unrelated to any

subsequent zircon-forming process.

140 Cretaceous

5.5 Cretaceous-Aged Zircon Overgrowths

As briefly mentioned in Chapter 4, and described in the results section of that

chapter, most of the Cycladic Series rocks contain zircons which have Cretaceous-aged

rims that may be interpreted as metamorphic. Determining whether these overgrowths are

metamorphic or magmatic and the significance of their age are discussed in this section.

The Cretaceous-aged zircon overgrowths apparent in many of the Series rocks of the

Cyclades show many characteristic features. They occur as distinct overgrowths, as

opposed to recrystallised zones, and Cretaceous ages are never found in the cores of

oscillatory-zoned grains and are only rarely found in the cores of irregularly zoned grains

(Figure 5-8).

50 µm

NX9451a.

200 µm

IO9615a. b. FL9602

50 µm

2880

91138

grain 9

grain 10

50 µm

NX94121

grain 25b

d.

127

c.

grain 23b 50 µm

NX94121

97

Figure 5-8: Cretaceous-aged metamorphic overgrowths from Cycladic Series rocks.: a) IO9615, darkunzoned rims surrounding oscillatory-zoned cores; b) grain 9, late Cretaceous overgrowth surroundingoscillatory-zoned early Cretaceous core and grain 10, Neoproterozoic oscillatory zoned detrital grain; c)dark homogeneous overgrowth on oscillatory zoned grain; d) dark faintly zoned overgrowth truncating theboundary of a pre-existing oscillatory zoned grain.

The growths are generally unzoned, display low luminescence and have variable,

but generally low Th/U ratios. There are two possible interpretations to explain them: 1)

they represent new zircon growth in response to metamorphism; 2) they represent new

zircon growth during magmatic activity. As discussed briefly in Chapter 3, the formation

of new magmatic zircon in S-type granites typically occurs as rims and is often

Chapter 5 141preferentially concentrated on grain terminations where it may have a homogeneous

unzoned appearance (see Keay et al. paper in Appendix A). This type of new zircon

growth also occurs in response to partial melting and it is thus impossible to distinguish

new zircon growth related to partial melting processes during high grade metamorphism

from small amounts of magmatic zircon. However, the zircon overgrowths all occur in

metasediments, not granitic rocks, and homogeneous unzoned, low luminescent rims are

also characteristic of hydrothermally-precipitated zircon (Claoue-Long et al., 1995) and in

some cases even “recrystallised” or replaced zircon (cf. Pidgeon, 1992). For this reason,

the favoured interpretation for the origin of the Cretaceous-aged overgrowths in zircons

from the Series rocks is that they were produced in response to a metamorphic event.

The age distribution of the Cretaceous zircon overgrowths was analysed by

combining results from different samples. This was done due to the relatively small

number of overgrowths available from each sample. Some individual samples yielded

recognisable populations of Cretaceous-aged metamorphic zircon overgrowths, such as

those listed in Table 5-2. However, the easiest way to highlight the major Cretaceous-

aged zircon populations was found to be by combining samples, as in Figure 5-9.

Table 5-2: Cretaceous age populations found in individual Cycladic samples

Sample Island Rock-Type Age populations (Ma)

NX9121 Naxos calc-silicate 68.8 ± 0.8 (6)78.9 ± 1.5 (4)94.8 ± 0.8 (8)

121 ± 2 (6)138 ± 2 (4)

NX9463 Naxos calc-silicate 75.1 ± 0.8 (4)NX94106 Naxos pelite 122 ± 1 (3)NX9315 Naxos leucogneiss 97.3 ± 1.3 (4)89639 Ios blueschist 88.8 ± 1.8 (4)

108 ± 2 (3)136 ± 2 (3)

90346 Ios quartz-phengite schist 77.1 ± 0.7 (5)127 ± 2 (4)

A combination of mixture modelling and visual identification reveals the timing of

zircon development in samples from different areas of the Cyclades (Figure 5-9). The

Cretaceous rims have been divided into groups of analyses according to their occurrence

in samples from different islands; Sifnos, Ios or Naxos. While Cretaceous-aged rims

were found in zircons in Series rocks of other Cycladic islands, e.g., Sikinos, these were

not included due to the relatively small number of analyses compared to the results from

Ios, Sifnos and Naxos. Further subdivision of Naxos ages has been on the basis of their

location either inside the partially melted Naxos core (results described in Chapter 3) or

outside the Naxos core (results described in Chapter 4). Three distinct age populations

can be distinguished from the Naxos core by mixture modelling, at 100 ± 1 Ma (n = 10),

142 Cretaceous122 ± 1 Ma (n = 8) and 133 ± 2 Ma (n = 9) (Figure 5-9). In contrast, five distinct age

populations can be distinguished from the Naxos Series rocks at 69 ± 1 Ma (n = 5), 77 ±

1 Ma (n = 11), 95 ± 1 Ma (n = 11), 118 ± 1 Ma (n = 10) and 137 ± 1 Ma (n = 6) (Figure

5-9). The Ios Series rocks show one very distinct Cretaceous peak at 78 Ma (n = 8), with

smaller peaks at ca . 67, 105 and 128 Ma (Figure 5-9). It is apparent from Figure 5-9 that

all sample combinations show troughs in the age probability density curves at

approximately 120-110 Ma and 90-85 Ma, suggesting a similarity in this aspect of the age

distributions for all samples.

70 80 90 100 110 120 130 140

Naxos Core

0

1

2

3

4

5

6

7

8

Ios Series Rocks

Naxos non-core

Sifnos

Age (Ma)

No.

Ana

lyse

s

Combined Cretaceous Metamorphic Ages

Figure 5-9: Overlain age proability density curves for combinations of Series from different islands, toshow similarities and differences between islands, and within and outside the Naxos core.

Sample SIF9345 has very few Cretaceous-aged rims (although a large number of

Tertiary-aged rims will be discussed in Chapter 6) and no significant populations can be

distinguished for this time period (Figure 5-9). Forty-nine analyses were conducted on

the Sifnos sample which should reveal age populations which are approximately 10 %

abundant at the 95% confidence level. However, the abundance of metamorphic

overgrowths (which are typically small volumetrically) is probably quite low and thus

many age populations from growth rims may have been missed. More analyses,

preferably from other Sifnos zircon samples, would be required to make meaningful

inferences about the significance, or otherwise, of the lack of Cretaceous-aged

metamorphic zircon growths.

Chapter 5 143Combining all Cretaceous ages from all Series rocks (excluding the Syros ophiolite,

89642) allows some broad generalisations to be made about the time of formation of

Cretaceous metamorphic zircon overgrowths (Figure 5-10). Peaks in the age ranges 80-

70 Ma, 105-95Ma, and 140-120 Ma can clearly be identified, suggesting that these

periods were times of active tectonism when new zircon was developed. It is unclear what

sort of tectonic process may have led to the development of new zircon growth in these

samples, but the strong peak at 80-70 Ma corresponds to the timing of ophiolite formation

on Syros and also the timing of high-T metamorphism in the Cycladic Upper Unit.

0

2

4

6

8

10

60 80 100 120 140Age (Ma)

No.

of

Ana

lyse

s

Figure 5-10: Age probability density diagram overlain by a histogram with 2 Ma bin widthscombining 159 Cretaceous zircon ages from all Series rocks (excluding the Syros ophiolite, 89642) .

5.6 Discussion

5.6.1 Dating Ophiolite formation

The formation age of oceanic crust that has undergone eclogite-facies

metamorphism can be difficult to determine due to a lack of dateable primary minerals,

and to overprinting by metamorphic assemblages causing possible resetting of isotope

systematics (Bill et al., 1997). The presence of primary igneous zircon in an eclogite

sample from Syros (89642) enables U-Pb dating of the time of zircon growth and hence,

presumably, the time of magmatic crystallisation of the ophiolite protolith. The zircons of

89642 are typically magmatic in both appearance and Th/U content, as described in

Section 5.4.1. The recognition of magmatic versus metamorphic zircon is complicated,

as recognised in Section 5.4, but the metamorphic zircon in this study generally forms

distinct rims (Section 5.5 and Chapter 6). The zircons from the Syros ophiolite are

interpreted as representing the time of oceanic crust formation at ca . 75 Ma, which

144 Cretaceoussuggests the ophiolites contained in the Series rocks of the Cyclades are of Turkish

affinity.

5.6.2 Similarities between Upper Unit and Series rocks

A previously unresolved question in the Cyclades has been whether the Upper Unit

is correlative with the Series rocks or is exotic material (Patzak et al., 1994; Katzir et al.,

1996). The Cretaceous age for the Syros ophiolite along with Cretaceous metamorphic

ages found in the Cycladic Series rocks strongly suggests that the Series rocks are high-P

equivalents of the Upper Unit. The Upper Unit has not experienced the Alpine high-P

metamorphism of the Series rocks, but is comprised of a melange of granitoids,

amphibolite, gneisses, calc-silicates, marbles and ophiolites which have experienced high-

T metamorphism, and yield Late Cretaceous cooling ages (Dürr et al., 1978b; Reinecke et

al., 1982; Patzak et al., 1994). The Series rocks are comprised of a similar range of

lithologies that have experienced the Alpine high-P metamorphism but were not

considered to have undergone any earlier metamorphic events (Dürr, 1986). From the

SHRIMP U-Pb dating undertaken in this study, it is now known that the Series rocks are

mainly Triassic-Jurassic in age and have experienced Cretaceous, possibly high-T,

metamorphism. The Upper Unit consists of Cretaceous-aged ophiolites interlayered with

a range of presumably older lithologies which have experienced high-T Cretaceous

metamorphism, but not an Alpine high-P overprint (Reinecke et al., 1982). As the Upper

unit has not experienced the range of Alpine deformations seen by the Series rocks, the

fact that the two units are correlated means that the relatively uncomplicated Upper Unit

can be used as an analogue for studying the pre-collisional history of the Series rocks.

This discovery could potentially lead to a much better understanding of the character of

the Cycladic Series rocks than was previously possible.

5.6.3 Correlation between Cycladic High-P Metaophiolite units

Whether the Syros ophiolite is correlative with other ophiolites that have undergone

high-P metamorphism in the Cyclades, such as Sifnos and Tinos, cannot be verified by

age results at this stage but it is suggested by the geochemical affinities of these units.

The geochemistry of the ophiolitic melange on Syros is suggestive of a back-arc basin

environment of formation (Seck et al., 1996). The presence of units with bonninitic

affinities in the ophiolite sequences of Sifnos and Tinos (Bröcker, 1991; Mocek, 1996)

has also been explained in this manner. This back-arc setting is consistent with the

known geochemistries of the metabasites from the Upper Units of the Cyclades (Reinecke

et al., 1982; Patzak et al., 1994; Katzir et al., 1996) which show supra-subduction zone

affinities, adding weight to the arguments put forward in Section 5.6.2. This suggests

that an accretion-subduction zone was operating immediately prior to the Alpine

collisional event which produced high-P metamorphism. Whether this represented a

Chapter 5 145long-lived (~25 Ma) accretion-subduction process, as suggested by (Katzir et al., 1996) is

difficult to determine and will be discussed in the next section.

5.6.4 Constraints on the Timing of High-P metamorphism (M1)

The recognition of the Syros ophiolite as a fragment of Cretaceous-aged oceanic

crust provides a maximum age constraint on the timing of Alpine high-P metamorphism.

Previous geochronological data obtained from the eclogite-blueschist facies rocks of the

Series have only placed minimum constraints on the timing of M1 and could significantly

post-date the time of initiation of high-P metamorphism (Ridley, 1984b). The length of

time between ophiolite formation and metamorphism cannot be constrained with any

certainty, except that results from this chapter suggest that metamorphism occurred

between ca . 75 Ma (from the age of the Syros ophiolite) and ca . 55-45 Ma (Altherr et al.,

1979; Andriessen et al., 1979; Wijbrans and McDougall, 1988; Baldwin, 1996). The

Cretaceous-aged metamorphic zircon overgrowths reported in this Chapter would also

constrain the timing of high-P metamorphism. While the age of ophiolite formation is ca .

75 Ma, most of the Cretaceous-aged metamorphic zircon overgrowths pre-date this event.

If the Cretaceous metamorphic-zircon ages reflect convergent margin processes, then the

ophiolite formation, and presumably obduction, occurred after this date, which indicates

that high-P metamorphism could not have been initiated prior to 75 ma and that the

Cretaceous-aged zircons are not related to M1. An important observation is that these

Cretaceous-aged zircon overgrowths occur in Series sediments, mainly calc-silicates, and

are not found in any samples of the Basement orthogneisses. This suggests that the

Cretaceous metamorphic ages can not be correlated with the M0 assemblages reported for

the Cycladic basement in Chapter 3. For this reason the Cretaceous metamorphic

episodes are referred to as M?.

5.6.5 Correlation with Menderes Massif

The age of the Syros ophiolite presented in this Chapter strengthens arguments for a

direct correlation between the blueschists of the Cyclades and the Menderes Massif.

Strong similarities in lithology, age and geological history of the blueschist units of the

Cyclades and the Menderes Massif has recently been presented (Candan et al., 1997), and

was previously suggested by other workers (Dürr et al., 1978a; Okay, 1984; Okrusch et

al., 1984; Ridley, 1984a). These similarities are in direct contrast to the differing ages of

Basement units in both areas, which are Devonian-Carboniferous and Precambrian,

respectively, as discussed in Chapter 3. Like the Cyclades, the Menderes Massif consists

of metamorphic core complexes where a cover series of interbedded Mesozoic marble-

schist sequences structurally overlie an older core or basement comprised largely of

orthogneisses (Dürr et al., 1978a; Bozkurt and Park, 1994; Hetzel and Reischmann,

1996). As pointed out by Onay (1949) and quoted in (Dürr et al., 1978a), the cover

series contains a distinctive metabauxite-bearing marble sequence (Dürr et al., 1978a)

146 Cretaceousremarkably similar to that found in the Series rocks of the Cyclades (Feenstra, 1985).

The protoliths of the Menderes Massif have been identified as Cretaceous in age (from

well-preserved rudist fossils in marble) and the meta-olistostrome sequence preserved in

the blueschists has been likened to that of Syros (Candan et al., 1997). Sequences from

both areas have undergone a Tertiary high-P, low-T event which was subsequently

overprinted by a Barrovian-type medium-P metamorphism under greenschist to

amphibolite facies conditions (Altherr et al., 1979; Sengör, 1984; Satir and Friedrichsen,

1986; Candan et al., 1997). Ridley (1984a) suggested that the Cycladic blueschists are

linked to convergence of the Menderes Massif and Sakarya continental blocks during the

Late Cretaceous-Early Tertiary and this is consistent with the age data presented in this

chapter.

5.6.6 Correlation with Pelagonian Zone, Internal Hellenides, Greece

The blueschist-bearing Series rocks of the Cyclades are thought to crop out

discontinuously through the Pelagonian zone, and are exposed in tectonic windows such

as Mt Olympos (Dürr et al., 1978a; Blake et al., 1981). Though there are many

stratigraphic similarities with the Cyclades, no Cretaceous-aged ophiolites have been

recognised from the Pelagonian zone and the ophiolites of the Pelagonian zone largely

seem to be structurally emplaced during the Late Jurassic-Early Cretaceous (Mountrakis,

1984). This relationship can also be seen in Evvia where a Triassic-Jurassic carbonate

platform of the Pelagonian zone was overriden by an ophiolite during the late Jurassic

(Robertson, 1991). This suggests that the convergent processes which resulted in

ophiolite emplacement on continental crust occurred much earlier in the Pelagonian zone

than in the Cyclades. This idea is also supported by the inferred timing of collision-

related high-P metamorphism in the Pelagonian. On the basis of 40Ar-39Ar ages from

phengites, it has been suggested that the Pelagonian zone has undergone two high-P

metamorphic events at ca . 100 Ma and 61-53 Ma (Schermer et al., 1990). Similar recent

work suggests that the first phase of high-P metamorphism may have occurred as early as

ca . 115 Ma (Lips et al., 1997). It is unclear whether the Cretaceous-aged metamorphic

overgrowths in the Cyclades can be related to such an early phase of high-P

metamorphism and no simple correlation between the Cyclades and the Pelagonian zone

can be made.

5.6.7 Correlation with External Hellenides

The Phyllite-Quartzite (PQ) unit exposed on Crete which was correlated to some of

the Series rocks described in Chapter 4 preserves evidence of high-P, low-T

metamorphism thought to be mid-Cretaceous in age (Seidel, 1977). This is in keeping

with estimates of high-P metamorphism from the Pelagonian zone and with the timing of

metamorphism of unknown character which produced new zircon growth in the

Chapter 5 147Cyclades, as decribed in the previous section. The ophiolites preserved in Crete are much

older than those from the Cyclades (Mid-Jurassic as opposed to Late Cretaceous) and

seem to have “Greek” rather than “Turkish” affinities (Koepke et al., 1985). As for the

Pelagonian zone, no simple correlation between the Cyclades and the external Hellenides

seems to exist, however Late Cretaceous ages have been reported for ophiolites from the

Subpelagonian - an intermediate zone between the Pelagonian and external Hellenides -

which has been taken as evidence that a Tethys ocean existed in some form until at least

this time (Clift and Robertson, 1989). The Syros ophiolite might also have formed a

fragment of this ocean.

5.6.8 Tectonic Implications

The recognition of a Cretaceous-aged ophiolite in the Cyclades extends the

geographic range of “Turkish” ophiolites and should prove useful in tectonic

reconstruction of the area. The existence of late Mesozoic/early Cainozoic subduction

zones in the Aegean has been suggested on the basis of Cretaceous metamorphic ages

obtained from phengites (Blake et al., 1981). This was despite the fact that most of the

ophiolites in adjacent mainland Greece were early Mesozoic in age (Hynes, 1974;

Roddick et al., 1979). Since this time, Cretaceous or “Turkish”-aged ophiolites have

been recognised in Greece in the Ermioni Complex of the Peloponessus (Clift and

Robertson, 1989), part of the sub-Pelagonian zone located between the external

Hellenides and the Pelagonian zone. These can be related to the operation and eventual

collapse of spreading centres related to convergence of the Eurasian and African

continents (Figure 5-11). During the Late Jurassic, complex triple junctions existed

throughout the Aegean and the motion of fragments of both continental and oceanic crust

was unconnected, or only loosely connected, to that of the two major plates (Africa and

Eurasia) (Burchfiel, 1980). In general, “Greek” ophiolites were mainly emplaced soon

after formation, during closure of ocean basins in Late Jurassic-Early Cretaceous time

(Mountrakis, 1986) (Figure 5-11).

148 Cretaceous

Moesia

Rhodope/Serbo-Macedonia

Kirsehir

Sakarya

MenderesTauride

E. Tauride

PuturgeBitlis

South Aegean

PindosOcean

Pontide-Caucasus Volcanic Arc

EURASIA

AFRICA

Alanya


Pelagonian

EARLY CRETACEOUS ~ 120 Ma

ophiolite obduction

Figure 5-11: Plate reconstruction for the early Cretaceous showing the location of the Cycladesas part of the South Aegean block, still in close proximity to north Africa (Robertson and Dixon, 1984).

The plate reconstruction shown in Figure 5-11 suggests that spreading ridges

developed behind the South Aegean block (which equates with the position of the

Cyclades) separating it from north Africa. The early Cretaceous obduction of “Greek”

ophiolites is shown along the margins of the Pelagonian zone.

Chapter 5 149During the Late Cretaceous, spreading ridges that separated the continental blocks

of “Turkish” affinity from each other and from the northern margins of Africa and Arabia

all came under active compression and collapsed, producing ophiolites (Robertson and

Dixon, 1984) (Figure 5-12).

Kirsehir

Menderes/Tauride

E. Tauride

Puturge

Bitlis

South Aegean

EURASIA

Alanya

Pelagonian

LATE CRETACEOUS

Iran

Sakarya

Hatay &Baer-Bassit

Black Sea

~ 80 Ma

obduction ofophiolites

AFRICARobertson and Dixon (1984)

Figure 5-12: Plate reconstruction for the eastern Mediterranean for the Late Cretaceous (fromRobertson and Dixon, 1984).

The plate reconstruction in Figure 5-12 shows little alteration in the position of the

South Aegean block (where the Cyclades would be located), from the Early Cretaceous

reconstruction, but does show widespread obduction of “Turkish” ophiolites onto the

Menderes/Tauride, Alanya, Puturge and Bitlis continental blocks and also onto the

margins of Arabia. Ophiolite obduction in the South Aegean block would presumably

have occurred some time after 75 Ma, but prior to high-P metamorphism at ca . 50 Ma.

150 Cretaceous5.7 Synthesis

Dating of an eclogite facies metaophiolite from the island of Syros, which has not

been previously attempted, shows that the ophiolite was formed at ca . 75 Ma. This age

places an important upper age constraint on the timing of high-P metamorphism in the

Cyclades. Most of the metasedimentary Series rocks preserve Cretaceous metamorphic

zircon ages which are possibly associated to the high-T metamorphism which affected the

Cycladic Upper Unit in the late Cretaceous and to high-P metamorphism reported in rocks

from the external Hellenides and the Pelagonain zone of Greece. The Syros ophiolite is

the same age and has geochemical similarities to the ophiolites described in the Cycladic

Upper units. This suggests that the Upper Unit provides an excellent analogue to the

Series prior to high-P metamorphism and thus can be useful in reconstructing the pre-

collisional history of these rocks. The age of the Syros ophiolite also indicates that it is of

“Turkish” rather than “Greek” affinity, suggesting a close relationship between the

Cyclades, the Menderes Massif and other Turkish continental blocks during the Late

Cretaceous. Such a relationship is supported by close similarities in the age, stratigraphy

and geological history of the Series rocks of the Cyclades and the blueschists of the

Menderes Massif, as recognised by (Candan et al., 1997). This correlation supports the

plate tectonic reconstruction of (Robertson and Dixon, 1984) for the Late Cretaceous

which has the South Aegean block in close proximity to the Turkish blocks (Figure 5-12).

Robertson and Dixon (1984) recognise this period as one of ophiolite obduction for the

“Turkish” crustal blocks, associated with collapse of oceanic spreading ridges during

convergence of the African and Eurasian plates. There are several problems with

correlations between the Series rocks of the Cyclades and the Pelagonian zone or External

Hellenides, but the intermediate Subpelagonian zone also preserves “Turkish”-aged

ophiolites, similar to those of the Cyclades.

Chapter 6 151

6. TERTIARY METAMORPHIC EVOLUTION OF THE CYCLADES

6.1 Introduction

The Tertiary tectonic evolution of the Cyclades includes polyphase metamorphism

and deformation including several episodes of fluid infiltration, anatexis and shearing.

Hence it provides a natural laboratory to assess the behaviour of a range of accessory

phases developed in response to differing metamorphic processes. For this reason the

following chapter tackles two main points, the timing of metamorphic mineral growth in

the Cyclades and its relation to tectonic processes, and the influence of fluid infiltration

associated with metamorphism. Both U-Pb isotope results and stable isotope work is

discussed and the results integrated into a general discussion used to construct a schematic

P-T-t path describing the Tertiary metamorphic evolution of the Cyclades.

The metamorphic evolution of the Cyclades is known to be Tertiary from previous

geochronological studies applying K-Ar, Rb-Sr and 40Ar-39Ar dating techniques (Altherr

et al., 1979; Andriessen et al., 1979; Maluski et al., 1987; Wijbrans and McDougall,

1988; Wijbrans et al., 1990; Baldwin, 1996). Two main metamorphic events are

recognised in the Basement and Series rocks exposed on different Cycladic islands: M1 a

high-P low-T event and M2 a medium-P, medium to high-T overprint (Buick and

Holland, 1989). M1 is thought to be related to Eocene collision of the Turkish and

Eurasian plates, while M2 has been variously described as Oligocene or Miocene

depending on the island and its particular metamorphic history.

There is increasing evidence that M2 actually represents several distinct

metamorphic episodes of similar grade in the time interval 30-19 Ma in the Cyclades

(Wijbrans et al., 1990). M2 has been subdivided into M2a, M2b and M2c (Buick, 1991)

where M2a is thought to be greenschist facies metamorphism of regional extent, M2b is

localised upper amphibolite facies metasomatism with associated partial melting on Naxos

(Jansen and Schuiling, 1976; Buick and Holland, 1989) and M2c is retrogression

associated with uplift (see Buick, 1991). M2b is thought to have occurred in the interval

26-20 Ma, most probably at the younger (20-19 Ma) end of this interval (Wijbrans and

McDougall, 1988; Buick, 1991). On Sifnos and Tinos there is evidence for at least three

periods of metamorphism, although it has been argued that these result from differential

uplift of segments of Sifnos that have only undergone M1 and M2 but that exposure of

different crustal levels makes it appear that there are are more than two metamorphic

episodes (Grütter, 1993; Wijbrans et al., 1993). On Tinos several retrograde reactions

are recognised as taking place after high-P metamorphism from 32-28 Ma either indicating

a retrograde period after blueschist metamorphism caused by decompression during uplift

or, alternatively a separate intra-Oligocene stage of prograde greenschist metamorphism

(Bröcker et al., 1993). The M2 overprint of M1 blueschist assemblages has been ascribed

to fluid flow (Bröcker, 1990).

152 Tertiary


Until isotopic studies of the Cyclades were undertaken (e.g., Andriessen, 1978),

little was known about the temporal evolution of the region except through rare

biostratigraphic markers and scattered correlations with other Alpine terranes. To date,

most of the geochronology in the Cyclades recorded the Tertiary history of the area. This

is a product of the isotopic work conducted: K-Ar, Rb-Sr and 40Ar-39Ar studies of

minerals with relatively low Tc as opposed to U-Pb investigations of higher temperature

mineral phases which might record pre-Tertiary events. Furthermore most pre-1990

isotopic measurements were of mineral separates rather than individual mineral grains

(Wijbrans et al., 1990), so that within-grain variations cannot be assessed. The lack of

information about spatial variations in isotope compositions within analysed mineral

grains makes these ages difficult to interpret. For example, 40Ar-39Ar ages may be

affected by mixing of several different generations of mineral growth (Wijbrans and

McDougall, 1986), excess argon (Altherr et al., 1979; Grütter, 1993) such as that

introduced along grain boundaries by fluids (Reddy et al., 1996) and/or retention of

“inherited” radiogenic argon in minerals (Monie et al., 1991; Lister and Baldwin, 1996;

Hames and Cheney, 1997; Kelley et al., 1997). These various processes may all result in

similar argon spectra (e.g., Lister and Baldwin, 1996; Scaillet, 1996) leading to problems

in the interpretation of 40Ar-39Ar spectra and casting doubt on the derived ages.

In the following section the previously obtained geochronological constraints on the

Tertiary metamorphic evolution of the Cyclades are described for each island investigated

in this chapter, in turn.

6.2.1 Naxos

(Andriessen, 1978; Andriessen et al., 1979) provided a comprehensive

geochronological framework for the geology of Naxos using extensive K-Ar and Rb-Sr

age dating. Four main periods of metamorphism were distinguished by these techniques:

M1: an Alpine high-P/low-T event at 45 ± 5 Ma (Middle Eocene);

M2: a medium-P/low to high-T Barrovian event dated as 25 ± 5 Ma (Late Oligocene/Early

Miocene);

M3: a contact metamorphic event initiated by the intrusion of a post-M2 granodiorite with

a minimum age of 11.1 ± 0.7 Ma determined from a Rb-Sr whole rock isochron derived

from aplitic/pegmatitic dykes cutting the main granodiorite body;

M4: a possible metamorphic event at 10 Ma associated with brittle deformation,

constrained by K-Ar dating of a pseudotachlyte veinlet, and also corroborated by four

mineral dates from other areas of Naxos.

Chapter 6 153

An apparent progression from old ages in the south-east where M1 mineral assemblages

are best preserved, to young ages in the migmatite core of Naxos has been observed

(Andriessen et al., 1979) (Figure 6-1).

0 2 4 6km

NW SE

1000500

20

40

60biotite hornblende phengitemuscovite

0

Sediment Upper Unit Marble Schist Amphibolite Leucogneiss Core

Age

(M

a)m

etre

s

Figure 6-1: NW-SE cross-section through Naxos showing the distribution of argon ages and anapparent decrease in ages towards the core of the island (adapted from Andriessen (1978) and using datafrom Wijbrans and McDougall (1988)).

More recent dating on Naxos by 40Ar-39Ar suggests that there is a progression to

younger ages towards the core of the island. This distribution has been suggested to

represent mixing of at least two separate white mica populations in the multigrain analysis

(Wijbrans and McDougall, 1986). Step-heating analysis of these concentrates yielded

generally upward convex age spectra for all but the samples from the core of Naxos,

where the M2 overprint was strongest. In the core, relatively flat age spectra were

obtained suggesting rapid cooling of the muscovite through its Tc for argon (~350 ˚C).

The convex upward patterns were interpreted as the result of mixing of two distinct white

mica populations: M1 high-P phengitic muscovite and M2 high-T muscovite. Maximum

white mica argon ages decrease as the core is approached and range from 43.3 Ma in

Zone 1 to 17.2 Ma in Zone 4. The apparent ages in the initial and final steps are variable

but generally lower in initial steps than final steps producing a convex-upward spectra.

Step heating experiments at high-T showed that argon release from pure M2 muscovite

lags slightly behind that of M1 phengite at temperatures in excess of 900 ˚C, suggesting

differences in the physical properties of these micas (Wijbrans and McDougall, 1986).

These convex upward age spectra have been interpreted as indicating that phengites

present in Zone 4 of Naxos where the stable M2 mineral assemblage includes muscovite,

biotite, garnet and staurolite still retain some radiogenic argon which, from diffusion

theory, could only occur if M2 was of short duration (Wijbrans and McDougall, 1986).

Other interpretations of convex upward age spectra have been made in studies

elsewhere. For example, Scaillet et al. (1992) reported an increase in Ar retention with

increasing Mg content in phengites that is recorded by characteristic bulk-sample convex

154 Tertiary

upward age spectra and even convex upward single-grain age spectra. This isotopic

discordance is attributed to compositional changes during re-equilibration of high-P

phengites during retrogression.

Ignoring the possible difficulties in interpreting argon results from high-P micas,

combined 40Ar-39Ar dating of hornblende and micas from Naxos (Wijbrans and

McDougall, 1988) have been used to suggest that M1 occurred before 50 Ma, the age of

white micas from Zone 1, while ages for M2 varied according to metamorphic grade with

older hornblende ages recorded in lower grade zones, ranging from 19.8 ± 0.1 to 15.0 ±

0.1 Ma. Hornblendes from a high grade section of northern Naxos yielded plateau ages of

11.6 ± 0.1 Ma and 13.8 ± 0.1 Ma area (Wijbrans and McDougall, 1988). These younger

ages could reflect thermal disturbance of argon spectra by the intrusion of numerous S-

and I-type granitoids in this part of Naxos (which is consistent with age constraints from

these granitoids in Chapter 7). The most accurate estimate for the timing of peak M2

metamorphism is considered to come from the age spectrum of a hornblende from a

sample collected at the M2 chloritoid-out isograd, which probably formed at less than 500

˚C, so that its plateau age of 16.1 Ma may record the time the hornblende crystallised as

the Tc for Ar in hornblende is ~ 525 ˚C (Wijbrans and McDougall, 1988). Younger ages

of 11.8 ± 0.1 Ma for muscovite and 11.4 ± 0.1 Ma for biotite from the migmatite core of

Naxos are regarded as representing rapid cooling from peak M2.

It is possible that more than one M2 Barrovian metamorphic episode occurred, as

evidenced by ages of 30-19 Ma found using single crystal argon geochronometry

(Wijbrans et al., 1990). Since this observation was made, the M2 event on Naxos has

been subdivided into three separate events, M2a (23 Ma) representing the timing of

regional Barrovian greenschist facies metamorphism (dated by Rb-Sr and K-Ar of low

grade muscovite), M2b (20-19 Ma) the timing of a higher grade thermal overprint

(Andriessen and Jansen, 1990) and M2c a retrograde overprint thought to be recorded by40Ar-39Ar amphibole ages of 16-15 Ma from the core reported by (Wijbrans and

McDougall, 1988; Buick, 1991).

6.2.2 Sifnos

To constrain the timing of high-P metamorphism and later greenschist retrogression

on Sifnos, several different methods have been used. Altherr (1977) and Altherr et al.

(1979) report concordant Rb-Sr and K-Ar ages of ca . 42 Ma from phengites from

northern Sifnos that they interpret as representing the end of high-P metamorphism. The

age for high-P metamorphism on Sifnos is in accord with results from other Cycladic

islands and from southeastern Naxos (Altherr et al., 1979; Andriessen et al., 1979).

Phengites from samples that show some evidence of overprinting, yielded ages of 48 and

41 Ma with consistently lower Rb-Sr ages (37-33 Ma). These rocks have been

overprinted by the later M2a greenschist facies metamorphism and so the discrepancy

Chapter 6 155

between K-Ar and Rb-Sr ages was interpreted as being due to the presence of excess

argon. Excess argon is found in amphiboles from northern Sifnos whose ages increase

with decreasing K content (age range 480-130 Ma) and is also evident in low-K minerals

such as chlorite (Altherr et al., 1979) so its influence must be considered. Altherr et al.

(1979) also report K-Ar and Rb-Sr ages on phengite from central Sifnos that vary

between 24-21 Ma and are thought to constrain the timing of M2 greenschist facies

metamorphism. A titanite fission-track age of 15.1 ± 2.1 Ma that is interpreted as the age

of cooling after the M2 Barrovian overprint (Wagner in Altherr et al., 1979).

More recent dating work utilising laser 40Ar-39Ar ages from individual mineral

grains (Wijbrans et al., 1990) yielded ages ranging from 42-36 Ma. Samples were taken

from Sifnos’s eclogite-blueschist domain (EBD), the main marble unit and the underlying

greenschist domain (GSD). Modelling of these data by (Lister and Baldwin, 1996) was

used to suggest that 42-40 Ma ages from the EBD represented cooling after M1 while

recrystallised phengites from retrograde shear zones that yielded ages in the range 34-28

Ma were interpreted to represent cooling during uplift, hence constraining the timing of

greenschist metamorphism on Sifnos to pre-30 Ma.

6.2.3 Ios

An estimate for the timing of M1 on Ios come from the 40Ar-39Ar spectra of a fresh

blueschist sample which contains evidence for closure at ca .39 Ma thought to reflect peak

or post-peak M1 followed by partial resetting at ca . 29 Ma in response to rehydration

during decompression (Grütter, 1993). Several generations of white mica have been

identified on Ios (Henjes-Kunst and Kreuzer, 1982; Baldwin and Lister, 1998). The first

is thought to preserve ages associated with M0 Variscan amphibolite metamorphism

affecting the basement of Ios ca . 500-300 Ma. Second generation micas yield K-Ar ages

of 39-34 Ma which were related to M1 by Henjes-Kunst and Kreuzer (1982) while

Baldwin and Lister (1998) report older 40Ar-39Ar apparent ages ranging from ca . 58-42

Ma with a plateau at 54 Ma thought to approximate the timing of M1. A third generation

of sericitic mica thought to be associated with M2 greenschist facies metamorphism yields

a K-Ar age of 25.7 Ma (Henjes-Kunst and Kreuzer, 1982), while 40Ar-39Ar apparent ages

of ca . 32-31 Ma and also ca . 21 Ma are thought to reflect recrystallisation under

greenschist conditions (Baldwin and Lister, 1998) and so the exact timing of M2 is

unclear. Grütter (1993) argues that the timing of M2 on Ios and most of the Cycladic

islands will be impossible to define using argon systematics because of the introduction of

excess argon presumably via M2 retrogressive fluid. Grütter (1993) considers the 30-19

Ma 40Ar-39Ar ages found in retrogressed samples from Syros, Sifnos and Tinos represent

maximum ages until the possibility of excess argon is eliminated and suggests that the

youngest ages of M2 reported for the Cyclades (e.g., Naxos) should be considered as

maximum M2 ages which would constrain M2 to be younger than 15 Ma (Wijbrans and

McDougall, 1988).

156 Tertiary

A range of possible contact metamorphic (M3) ages associated with granite

intrusions are reported from K-Ar on hornblende and micas from Paros, yielding a range

from ca . 14-10 Ma (Altherr et al., 1982). 40Ar-39Ar thermochronology on some M1 K-

feldspars from the Ios basement reveals argon loss during a ca . 14 Ma event thought to be

associated with magmatic activity (Baldwin and Lister, 1998) and also supported by a

whole rock-phengite Rb-Sr age of ca . 13 Ma from a meta-aplite dyke from the basement

(Henjes-Kunst and Kreuzer, 1982).

6.2.4 Other Cycladic Islands

The timing of M1 metamorphism on Syros which, like Sifnos, contains well

preserved M1 assemblages has been dated using argon systematics (Maluski et al., 1987;

Baldwin, 1996). White mica concentrates from samples which have been variously

overprinted to greenschist facies yield ages ranging from 53-36 Ma (Maluski et al., 1987;

Baldwin, 1996). Maluski suggests high-P metamorphism occurred between 117-53 Ma,

the upper bound defined by a K-Ar age on glaucophane from Syros which may have been

affected by excess argon, while Baldwin (1996) restricts M1 to between 54-50 Ma

according to age results from unretrogressed samples. Three groups of 40Ar-39Ar ages

have been distinguished from samples from Tinos (Bröcker et al., 1993), with high-P M1

metamorphic assemblages yielding ages of 45-39 Ma, and M2 samples yielding spectra

with ages of 30 Ma and upward convex spectra with ages of 23-21 Ma. All argon spectra

reported by Brocker et al. (1993) are disturbed, suggesting the Tinos samples were

affected by at least two periods of disturbance which have been interpreted according to

three periods of Tertiary metamorphism, M1 high-P metamorphism with a minimum age

of 44-40 Ma, overprinting greenschist metamorphism during exhumation between 32-28

Ma followed by a later greenschist overprint with a maximum age of 23-21 Ma,

interpreted to be the product of mixed ages. These results for Tinos are very similar to

those reported by Wijbrans et al. (1990) for Sifnos, where a second period of greenschist

metamorphism at ca . 19 Ma is also recognised. The timing of an earlier phase of

greenschist metamorphism is supported by a K-Ar age from white mica reported from a

greenschist on Milos at ca . 33 Ma (Fytikas et al., 1976). The prceding compilation of

data indicates that there is good evidence for at least two periods of overprinting

greenschist metamorphism on most of the Cycladic islands.

6.3 Evidence of Fluid Infiltration in Cycladic Rocks

6.3.1 Naxos

The nature of fluid-rock interaction on Naxos has been the focus of considerable

work. An important debate concerning the significance of CO2-rich fluids in

metamorphism was centred on Naxos following the publication of a paper by Schuiling

and Kruelen (1979) suggesting that the amphibolite facies M2b overprint on Naxos

Chapter 6 157

occurred in response to heating by CO2-rich fluids derived from the mantle. This idea

was proposed to explain several observations that had been made on Naxos:

1) the high XCO2 contents in fluid inclusions in many different rock-types

2) the mantle-like δ13 C values of these fluid inclusions, in contrast to the “normal” δ13

C found in metamorphic schists on the island

3) stable isotope evidence for large-scale isotopic exchange in the metamorphic complex

(Rye et al., 1976).

Kruelen (1980) found that individual CO2-rich fluid inclusions had relatively low

δ13C (-1 to -5%), thought to be consistent with the gas being derived from a deep-seated

source. Modelling the thermal effects of flushing CO2-rich mantle fluids through

sediments suggested that metamorphism could be the result of an input of advective heat

caused by a mantle CO2 flux (Schuiling and Kreulen, 1979). Further modelling by Bickle

and McKenzie (1987), suggested that such a mechanism would need to operate at

relatively fast rates over ~ 0.1 Ma to achieve the necessary patterns of heat distribution.

As the low porosity of marbles on Naxos precludes them from significant fluid infiltration

over this short time scale, Jansen et al. (1989) argued against this model.

Detailed observations of carbon and oxygen isotope profiles across the contacts

between schist and marble layers on Naxos suggested that the marble was altered over a

small-scale (metres) by fluids derived from dehydration of the surrounding pelites during

prograde M2b metamorphism (Bickle and Baker, 1990). However, no evidence of

pervasive CO2-rich fluids operating during the Barrovian metamorphism was found

(Baker et al., 1989). These observations seem inconsistent with the ubiquitously high

CO2 contents of fluid inclusions in all samples from Naxos, unless the inclusions are not

representative of peak metamorphic fluid compositions. The possibility that these

chemically distinct fluid inclusions might result from selective water leakage from mixed

H2O-CO2 inclusions was suggested by Buick and Holland (1991): a process that has been

experimentally verified by Hall and Sterner (1993) and Bakker and Jansen (1994). The

recognition of post-metamorphic regionally penetrative low δ13C calcite veins (Ganor et

al., 1994) may also explain the low δ13C found by Kreulen (1988). These recent studies

suggest that, while fluid infiltration has been an important process on Naxos, the

composition of the fluid is not that of mantle-derived CO2-rich fluids and is more likely

meteoric H2O (Lewis and Holness, 1997).

6.3.2 Sifnos

Sifnos, unlike Naxos, has not undergone a high-grade Barrovian metamorphic

overprint and preserves high-P metamorphic assemblages formed at ~ 450-500 ˚C and ~

15 kbars in the northern part of the island (Schliestedt and Okrusch, 1988), overprinted

by greenschist grade assemblages in the southern half of the island. The preservation of

M1 assemblages in northern Sifnos has been ascribed to restricted fluid infiltration

(Matthews and Schliestedt, 1984). A large marble unit separating dominantly blueschist

158 Tertiary

mineralogies from greenschist ones was thought to have acted as a barrier to the passage

of fluids, although recent structural work suggests that units may be tectonically

juxtaposed (Avigad et al., 1992). Oxygen isotope evidence suggests the greenschist

overprint on Sifnos was governed by infiltration of 18O-CO2 enriched aqueous solutions

(Matthews and Schliestedt, 1984). Isotope profiles across lithologies suggest that

infiltration of fluids was not pervasive and was limited to localised flow in individual

layers (Putlitz et al., 1994).

6.3.3 Ios

As recognised for Sifnos, externally-derived hydrous fluids caused the

development of widespread overprinting M2 metamorphic mineral assemblages during

decompression of the metamorphic complex on Ios (Grütter, 1993). The introduction of

these water-rich fluids (XCO2 < 0.06) is thought to occur at the onset of, and perhaps as a

result of D2 deformation (Grütter, 1993) associated with extension and exhumation of the

Ios metamorphic core complex.

6.3.4 Other Cycladic Islands

The presence of interlayered blueschists and greenschists on Tinos allowed an

investigation of whether retrogression is controlled solely by infiltration of a fluid phase

or by compositional differences between various rock-types (Bröcker, 1990). Stable

isotope and textural studies by Brocker (1990; 1993) showed that fluid-to-rock ratios

were low and that the extent of the greenschist overprint was controlled by availability of

a fluid phase, which was preferentially channelised into specific structures or along

contacts between rocks of different permeability.

6.4 Compilation of Events and Age Data for the Cyclades

Geochronological information from the various rock units of the Cyclades can be

used to construct a probable sequence of events that have affected the region since the

Cretaceous, and this is shown in schematic form in Figure 6-2A. Periods where ingress

of fluids is likely to be a factor in the metamorphic evolution of the rocks have been

denoted by a star. Generalised P-T paths for the main metamorphic episodes in the

Cyclades are illustrated in Figure 6-3A. In Chapter 5, the age of the Syros ophiolite was

constrained at ca . 75 Ma, consistent with age determinations for other ophiolites of

“Turkish” affinity in the region. This age is taken to reflect the timing of oceanic crust

formation. The development of metamorphic soles up to amphibolite-facies grade in

association with ophiolites has long been recognised (Williams and Smyth, 1973;

Woodcock and Robertson, 1977). Such metamorphism in the eastern Mediterranean

occurs closely after ophiolite formation, and is thought to record the timing of tectonic

slicing of oceanic lithosphere which precedes obduction of the ophiolites onto continental

crust by 25-30 Ma (Spray and Roddick, 1980; Thuizat et al., 1981; Whitechurch et al.,

Chapter 6 159

1984). Evidence of high-T metamorphism is preserved in the Upper Unit of the

Cyclades, which essentially consists of an ophiolitic melange, and which has not

experienced M1 or M2 (Reinecke et al., 1982; Patzak et al., 1994). The relationship

between the timing of metamorphism in the Upper Unit and whether it can be related to

ophiolite formation and tectonic slicing is unclear. However, if, as is argued in Chapter 5,

the Upper Unit represents a good analogue for the Series rocks of the Cyclades, then it

can be inferred that the Series rocks have also experienced this high-T metamorphism

which will be denoted M? to avoid confusion with the M0 described from Basement rocks

of the Cyclades in Chapter 3. This was closely followed by a retrograde overprinting

metamorphic event recorded by the rejuvenation of argon ages in micas to 63 and 55 Ma

(Altherr et al., 1979; Patzak et al., 1994) (Figure 6-2).

Ophiolite formation was most likely accompanied by sea-floor alteration. It has

been argued, from stable isotopic evidence of high δ18O values, that the protoliths to the

high-P metabasic rocks of the Cyclades were extensively affected by sea-floor alteration

(Matthews and Schliestedt, 1984; Putlitz et al., 1994) prior to Early Tertiary

subduction/collision. This alteration would provide a protolith that contains large

quantities of H2O in the form of chlorite and clay minerals which would devolatilise

during high-P metamorphism (e.g., Norris and Henley, 1976; Bebout and Barton, 1989;

Philippot and Selverstone, 1991; Nadeau et al., 1993). So two periods of early fluid

influx can be inferred, the first from externally-derived fluids (magma interaction with

sea-water) and the second associated with dewatering of this material during the

subduction/collision process.

160 Tertiary

D1

D2

D4

D3

75 Ma1Ophiolite Formation

Eclogite-facies Metamorphism54 - 50 Ma6,7

Greenschist-facies Metamorphism32-28 Ma10, 11,12

Retrogression

Fluid influx

Lower Amphibolite-facies Metamorphism16-15 Ma9

Contact Metamorphism

14-10 Ma8,9,14,15

Granite IntrusionShearing

Release of FluidsPartial Melting

Greenschist & Amphibolite-facies Metamorphism26-19 Ma5,9,10,11,13,14

RetrogressionFluid influx

Blueschist-facies Metamorphism

48 - 41 Ma5,8,9,10,11

Lawsonite breakdownRelease of H2O

Fluid influx

Sea-floor Alteration

High-T Metamorphism~70 Ma2,3,4

63 - 55 Ma3,5Greenschist-facies Metamorphism

Retrogression

Collisional Events

Release of FluidsPrograde Devolatilisation

Release of H2O

Clinozoisite BreakdownParagonite Breakdown

Tectonic slicing

ProgradeDevolatilisation

Figure 6-2: Schematic diagram illustrating inferred periods of metamorphism and fluid flow (denoted bystar). The timing of events has been constrained using data from the following sources: 1. This study; 2.Maluski et al. (1987); 3. Reinecke et al. (1982); 4. Patzak et al. (1994); 5. Altherr et al. (1979); 6.Baldwin (1996); 7. Baldwin and Lister (1998); 8. Andriessen et al. (1979); 9. Wijbrans and McDougall(1988); 10. Wijbrans et al. (1990); 11. Bröcker et al. (1993); 12. Lister and Baldwin (1996); 13.Andriessen (1991); 14. Andriessen and Jansen (1990); 15. Altherr et al. (1982).

Chapter 6 161

Conditions of high-P metamorphism in the Cyclades reached eclogite grade before

isothermal decompression during which many of the rocks retrogressed to blueschist

grade (collectively termed M1). The earliest estimates of the timing of high-P

metamorphism are 54-50 Ma from argon systematics, which probably reflect a minimum

age constraint on the timing of peak high-P conditions (see Ridley, 1984). The transition

from eclogite to blueschist-facies metamorphism is accompanied by release of H2O

associated with lawsonite breakdown. There is increasing evidence that subsequent

greenschist/amphibolite grade metamorphism (M2) in the Cyclades did not immediately

follow on from M1 and that the rocks cooled to ~ 300-250 ˚C prior to M2 which

corresponds to depths of 8 to 12 km (Wijbrans and McDougall, 1988; Wijbrans et al.,

1990; Wijbrans et al., 1993; Baldwin and Lister, 1998). During initial cooling from

peak-M1 conditions, greenschist-facies mineralogies were developed, possibly in

response to fluid influx from H2O released during the breakdown of high-P minerals such

as clinozoisite and paragonite. This would explain the timing of early greenschist-facies

metamorphism well before peak-M2 (Wijbrans et al., 1990).

ANDALUSITE

KYANITE

SILLIMANITE

P kb

ar

0

2

4

6

8

10

12

14

Temperature ˚C100 200 300 400 500 600 700

? ?

M1

M2M?

D1

D2

D3

D4

Figure 6-3: Schematic P-T path for Cycladic samples. Pathways for M1 and M2 adapted from P-Tpaths for Series rocks of Buick and Holland (1989), Avigad et al. (1992), and Wijbrans et al. (1993).Pathway for Cretaceous metamorphism constrained by the P-T estimates of Patzak et al. (1994) for theCycladic Upper Unit.

162 Tertiary

Peak M2 metamorphic conditions on Naxos were accompanied by partial melting in

the Naxos core and release of fluids during prograde devolatilisation (Buick and Holland,

1991) (Figure 6-2). A retrograde period of lower-amphibolite metamorphism was

probably associated with crystallisation of partial melts in the core (Buick and Holland,

1991), while fluid influx into surrounding Series rocks may have been aided by the

development of extensional shear structures post-peak M2 (Buick, 1991). The final

stages of fluid infiltration were related to granite intrusion and contact metamorphism

(M3). From the above discussion it is clear that interpreting the ages of metamorphic

minerals depends on an understanding of the complicated metamorphic history of the

Cyclades. It is possible to identify at least seven distinct periods when the rocks of the

Cyclades were influenced by internally or externally-derived fluids. If the U-Pb

accessory minerals dated during this study were produced by fluid infiltration, then a

range of metamorphic ages would be expected from these rocks.

6.4.1 Sample Selection

Most of the samples described in this section are from Naxos (locations shown in

Figure 6-4). In addition, two samples from Ios and a sample from Sifnos will also be

discussed. Naxos was selected for intensive study because it preserves an almost

complete sequence of Barrovian metamorphic isograds (M2) overprinting early M1 high-P

mineral assemblages (Jansen, 1973; Jansen and Schuiling, 1976). It was considered a

good place to test the applicability of SHRIMP U-Pb geochronology to rocks of various

metamorphic grade.

Chapter 6 163

1 532 km40

N

Granodiorite

Migmatite

Upper Unit, Alluvium

MarbleSchist

Fault

Age Contour

12Ma

15Ma

20Ma30Ma

40Ma

NX94120NX94121

NX9637NX9638

NX94103

NX9435

NX9319

NX9320NX9315

Figure 6-4: Naxos Sample Locations and contour map of argon ages from John and Howard (1995).

Naxos is a classic example of Barrovian style metamorphism, and of particular

interest because different workers have argued that a succession of short thermal pulses

might be involved (e.g., Wijbrans and McDougall, 1986; Wijbrans and McDougall, 1988;

Buick and Holland, 1989; Baldwin and Lister, 1998). If the study demonstrates that

periods of zircon growth coincide with these periods of elevated temperature it gives more

credence to this hypothesis. If there is a continuum of growth then this suggests long

lasting metamorphism as is currently accepted to be the case. The study then can place an

164 Tertiary

important constraint on the hypothesis that metamorphism has involved a sequence of

short-lived thermal pulses.

Zone 1 has experienced little influence from M2, Zone 2 contains some muscovite

and chlorite, Zone 3 shows crystallisation of biotite and almandine garnet, Zone 4 begins

with the reaction of chloritoid to staurolite in mica schists and metabauxites while kyanite

also appears, in Zone 5 kyanite and fibrolitic sillimanite coexist and kyanite was thought

to disappear in close proximity to the “migmatite” demonstrably core (although kyanite is

demonstrably present Buick, 1988) where peak metamorphic conditions have been

estimated as 6-7 ± 2 kbar and 670 ± 50 ˚C (Jansen and Schuiling, 1976; Buick and

Holland, 1989). The age results for zircons from all the samples discussed here has

already been examined in the preceding chapters concerned with protolith ages and it is

only the young (Tertiary) metamorphic zircon ages which are discussed in detail here. It

should be emphasised that the conditions of metamorphism in the Cyclades were not

expected to be conducive to the development of metamorphic zircon as M2 metamorphic

grades are relatively low (~650-700 ˚C in the Naxos core and < 640 ˚C outside the core).

In addition, ages from metamorphic monazite and titanite are also presented.

6.5 SHRIMP U-Th-Pb Zircon Results

6.5.1 Naxos

Distinct differences can be recognised in the morphology and age of new zircon

growth identified in rocks from within and outside the core of Naxos. These differences

may reflect the partial melting process which has affected rocks within the Naxos core but

not the rocks outside the core. For this reason the zircons from these two areas and their

ages will be discussed separately. All ages are split into groups according to whether they

are Miocene (and hence possibly related to M2) or early-middle Tertiary (and possibly

related to M1).

6.5.1.1 Samples from within the Naxos core

Several of the samples from the Naxos core, which are described in Chapter 3,

yielded Tertiary ages that have been interpreted as metamorphic. These samples include;

NX9315, NX9319, NX9320, NX94103, NX9638 and NX9637 (Appendix E). The

samples are all from the leucogneiss core of Naxos which Chapter 3 identified as

sedimentary Series rocks of Triassic age or younger, rather than Variscan basement as

suggested elsewhere (Andriessen et al., 1987). Rocks within the Naxos core provide

abundant field evidence of the operation of partial melting processes in the form of the

development of igneous textures (the majority of leucogneiss samples underwent H2O-

saturated melting at ~ the granite solidus (Buick and Holland, 1991), and the presence of

numerous pegmatitic pods, and the development of migmatites. The new zircon growth

Chapter 6 165

invariably occurs as homogeneous to faintly-zoned overgrowths with a low luminescence

commonly forming terminations or fine rims on pre-existing grains (Figure 6-5).

a. NX9319

100 µm

17

283

(grain 20)

100 µm

NX9637b.

17

22

grain 4

17 14

grain 17

19grain 14

50 µm

SIF9345c.

grain 26

54

50 µm

SIF9345d.

39295

grain 12

50 µmgrain 9

SIF9345

287

42

e.

50 µm

NX94120f.

57

163

227

33

grain 1

grain 2

h. NX94121

grain 24b 50 µm

37

50 µm

NX94121g.

grain 4b

18

54

Figure 6-5: Cathodoluminescence images of zircon. Figures a and b represent zircons from thepartially-molten core of Naxos, while the remaining figures depict Tertiary overgrowths on zircons fromoutside the core or from other islands which have not experienced partial melting: a) NX9319; b)NX9637; c) d) e) SIF9345; f) NX94120; g) h) NX94121.

166 Tertiary

The new zircon growth yields a Miocene population of ca . 18-17 Ma indicating that new

zircon growth occurred during a period of partial melting in the core (Figure 6-6). There

is an earlier period of zircon growth during the Miocene that can be identified from

mixture modelling and by visual inspection of Figure 6-6 at 22.8 ± 0.4 Ma (n = 5).

0

2

4

6

8

10

5 10 15 20 25

n = 62

No.

of

Ana

lyse

s

Age (Ma)

Zircon overgrowth agesNaxos Core

Figure 6-6: Miocene zircon ages from the Naxos core displayed on a combined age probability densitycurve overlain by a histogram of ages (with 0.5 Ma bin widths).

Palaeo-Eo-Oligocene aged zircon growths from the core can be identified at 28.7 ±

0.3 Ma (n = 12), 33.1 ± 0.6 (n = 4), 41.7 ± 0.4 Ma (n = 9) and 53.3 ± 0.5 Ma (n= 10)

(Figure 6-7). These ages are also common elsewhere in the Cyclades as shown below.

Chapter 6 167

0

1

2

3

4

5

6

7

25 30 35 40 45 50 55 60 65

n = 35N

o. o

f A

naly

ses

Age (Ma)

Zircon overgrowth agesNaxos Core

Figure 6-7: Palaeo-Eo-Oligocene zircon ages from Naxos core displayed on a combined age probabilitydensity curve overlain by a histogram of ages (with 2 Ma bin widths).

6.5.1.2 Samples from outside the Naxos core

New zircon growth can also be recognised in samples from Naxos outside the

leucogneiss core. It shows some morphological similarities to the zircon overgrowths

within the Naxos core, occurring mainly as low-luminescent rims that are generally

homogeneous and unzoned. However, it differs in being commonly irregular, spongy,

and inclusion-filled. These features are similar to those reported for zircon formed by

hydrothermal processes (Wayne and Sinha, 1992). Combining the ages of zircon growth

rims from outside the Naxos core yields several distinct age populations in the Miocene

(Figure 6-8) and Palaeo-Eo-Oligocene (Figure 6-9).

168 Tertiary

0

1

2

3

4

5

5 10 15 20 25

n = 15

No.

of

Ana

lyse

s

Age (Ma)

Zircon overgrowth agesOutside Naxos core

Figure 6-8: Miocene zircon ages from outside the Naxos core displayed on a combined age probabilitydensity curve overlain by a histogram of ages (with 1 Ma bin widths).

These Miocene zircon ages are all from one sample NX94121 that was exetsnively

studied and whose results are illustrated in Figure 6-10. A number of age populations can

be identified via mixture modelling at 13.9 ± 0.3 Ma (n = 2), 15.2 ± 0.3 Ma (n=5), 17.3

± 0.9 (n = 2) and 22.1 ± 0.8 Ma (n= 6). In the Palaeo-Eocene age populations occur at

33.6 ± 0.6 Ma (n = 19), 41.4 ± 0.7 Ma (n = 13), 47.7 ± 0.4 Ma (n = 17), 55.0 ± 0.6 Ma

(n = 6) and 63.4 ± 0.5 Ma (n = 6).

Chapter 6 169

0

1

2

3

4

5

6

7

25 30 35 40 45 50 55 60 65

n = 60

Zircon overgrowth agesOutside Naxos core

No.

of

Ana

lyse

s

Age (Ma)

Figure 6-9: Palaeo-Eo-Oligocene zircon ages from outside the Naxos core displayed on a combined ageprobability density curve overlain by a histogram of ages (with 1 Ma bin widths).

Most of these age populations can be distinguished in a single rock sample. This is

illustrated using sample NX94121, a calc-silcate which preserves evidence of multiple

episodes of zircon growth (Figure 6-10).

170 Tertiary

0

1

2

3

4

5

6

10 20 30 40 50 60 70

n = 74

No.

of

Ana

lyse

s

Age (Ma)

NX94121

Figure 6-10: Tertiary ages from zircons from sample NX94121 displayed on a combined ageprobability density curve overlain by a histogram of ages (with 1 Ma bin widths).

6.5.2 Sifnos

Unlike the zircon overgrowths described for Naxos, there are no Miocene zircon

growth ages are found in the Sifnos sample. Ages are concentrated in the Palaeo-Eo-

Oligocene, possibly reflecting the importance of this period in the metamorphic evolution

of Sifnos and are consistent with the observed lack of a high-grade M2 overprint on this

island.

Chapter 6 171

0

1

2

3

25 30 35 40 45 50 55 60 65Age (Ma)

No.

of

Ana

lyse

s

Zircon overgrowthsSIF9345

n = 16

Figure 6-11: Tertiary-aged zircon overgrowths from Sifnos, sample SIF9345, shown using a cumulatedprobability density curve overlain by a histogram with 2.5 Ma bin widths.

Ages identified visually and using mixture modelling show a small population at ca .

37.7 ± 1.1 Ma (n=3) and larger groups at 43.7 ± 0.9 Ma (n=5) and two separate peaks at

50.5 ± 1.1 Ma (n=4) and 54.9 ± 1.1 (n=4). The Sifnos zircons have experienced a

distinctly different growth history to those from Naxos or Ios with ages concentrated in

the range 35-69Ma. While Ios also lacks Miocene zircon growth, it does display zircon

growth during both the Palaeo-Eocene and Cretaceous whereas, as was shown in Chapter

5, the Sifnos zircons show little evidence of new growth during the Cretaceous.

Whatever the process responsible for promoting new zircon growth in this Sifnos sample,

it seems to be restricted to the Palaeo-Eo-Oligocene.

6.5.3 Ios

As for the sample from Sifnos, none of the Ios samples exhibit new zircon

development during the Miocene, but small Palaeo-Eo-Oligocene age populations are

found at 42 Ma (n=4) and 61 Ma (n = 7). In fact these ages are all common to zircons

from different samples from several islands (Table 6-1A) as shown Figure 6-12A, which

combines all Tertiary-aged metamorphic zircon ages. However, the Ios samples also

display significant new zircon growth in the Cretaceous (discussed in Chapter 5),

whereas only limited Cretaceous zircon was identified in the Sifnos sample.

172 Tertiary

0

1

2

3

25 30 35 40 45 50 55 60 65Age (Ma)

No.

of

Ana

lyse

sZircon overgrowthsIos Series rocks

n = 14

Figure 6-12: Tertiary zircon overgrowths from combined Ios Series rocks illustrated using an ageprobability density curve overlain by a histogram with 2.5 Ma bin widths.

6.5.4 Combined Metamorphic Zircon Age Results for Cyclades

No Miocene zircon ages were derived from the samples from Ios or Sifnos. The

combined histogram/age probability density curve (Figure 6-13) for zircon overgrowths

produced during the Miocene shows a distinct difference in the ages from the partially-

molten core of Naxos and the surrounding lower grade metasedimentary rocks. The

dominant zircon population from the core occurs at 18 Ma, while outside the core it

occurs at 14 Ma. The difference in Miocene ages is possibly a function of structural depth

during Miocene metamorphism. The Naxos core was at lower structural levels than those

units outside the core, and experienced higher temperature conditions as a result (Buick,

1991). The age distribution for samples from different structural levels is unexpected as

thermal modelling of crustal thickening predicts that deeper structural levels will show the

youngest ages (England and Thompson, 1984) whereas the opposite is observed on

Naxos. The higher temperature conditions in the core led to partial melting and formation

of leucogneiss bodies and the development of new zircon growth related to partial melting

processes. The zircons outside the core are demonstrably different, often consisting of

spongy, inclusion-filled overgrowths. Samples from outside the Naxos core did not

reach high enough grades to melt during the Miocene and the growth of zircons in these

rocks therefore reflects sub-solidus processes. Sub-solidus growth also holds true for

ALL the metamorphic zircon overgrowths older than ~ 18 Ma, the time of partial melting

in the Naxos core. This comparison shows that partial melting conditions are not required

for the production of metamorphic zircon growth.

Chapter 6 173

0

2

4

6

8

10

5 10 15 20 25

Combined Miocene Metamorphic Ages

Age (Ma)

No.

of

Ana

lyse

swithin Naxos core

outside Naxos core

Figure 6-13: Combined Miocene zircon ages from Naxos showing overlain cumulated probabilitydensity curves for samples from within the Naxos core and from outside.

0

1

2

3

4

5

6

7

25 30 35 40 45 50 55 60 65

within Naxos core

Ios Series

outside Naxos core

SifnosIos

Combined Palaeo-Eo-Oligocene Metamorphic Ages

Age (Ma)

No.

of

Ana

lyse

s

Figure 6-14: Combined Palaeo-Eo-Oligocene age distribution curves for zircons from Ios, Sifnos andfrom within and outside the Naxos core, overlain to identify age similarities and differences.

The Naxos core shows a major period of Palaeo-Eo-Oligocene zircon growth at 28

Ma, while zircon growth occurs earlier outside the core at 34 Ma. Both within and

outside the Naxos core an age population at ca . 42 Ma is identifiable, while a strong age

population at 48 Ma is evident outside the core but is not duplicated within the core.

174 Tertiary

There were also age differences in the timing of zircon growth from within and outside

the core during the Cretaceous, as mentioned in Chapter 5. New zircon growth from both

inside and outside the Naxos core and on Ios shows broadly the same age pattern,

suggesting that all these units experienced the same Cretaceous events leading to new

zircon formation. Figure 6-14 shows that all samples have an Eocene peak at 45-40 Ma,

which coincides with estimates of the timing of high-P metamorphism from other dating

techniques. There is another common Eocene-age peak at approximately 55 Ma. The

Sifnos zircons show some differences, although emphasis should not be placed on this

one sample because of the relatively small number of analyses involved.

The consistent presence of ca . 42 Ma ages from different samples from different

Cycaldic islands is interpreted here as the preferred estimate for the timing of M1 based on

U-Pb dating of zircons. These and other zircons are interpreted as being the product of

fluid infiltration, as there is no textural evidence of their formation from in situ mineral

breakdown reactions such as those described by Pan (1997), nor are they associated with

partial melting process. Morphologically, most of the Tertiary-aged zircons appear as

overgrowths so they do not have the appearance of recrystallised/replaced grains

(Pidgeon, 1992) (Figure 6-5). If the interpretation of these zircons as overgrowths

precipitated from infiltrating fluids is correct, then they have first formed (as opposed to

recrystallising) under high-P metamorphic conditions.

As shown in Chapter 3, Cycladic Basement rocks occur only in a thin tectonic

sliver located just outside the leucogneiss core of Naxos and they have not experienced

Miocene partial melting. Unlike other samples from outside the Naxos core, no Tertiary

metamorphic zircon growth is recognisable in the Basement rocks. Combining 179

analyses of orthogneiss and garnet-mica schist samples comprising the Variscan basement

of Naxos and the other Cycladic islands, which should detect any populations that are

greater than 0.03% abundant at the 95% confidence interval (Appendix D), no Tertiary

metamorphic ages are revealed. It therefore appears that, in contrast to the Series rocks

described in this section, no new metamorphic zircon developed in any of the rocks of the

Cycladic basement during the Tertiary, which suggests that lithology is playing a crucial

role in determining zircon growth. Most of the Basement rocks in the Cyclades occur at

the lowermost structural level and so the lack of zircon development suggests that

structural level and hence absolute temperature and metamorphic grade does not influence

whether new zircon growth occurs in the rocks of the Cyclades. This is in direct contrast

to studies of metamorphic zircon development in the contact aureoles of granite intrusions

(Davis et al., 1968; Ferry, 1996) which show an increase in zircon development up-

temperature that may be associated to fluid infiltration.

Chapter 6 175

Table 6-1: Summary of Tertiary age populations identified for zircon overgrowths fromindividual samples.

Sample Rock-Type Location Main Ages(No. Analyses)

NX9121 calc-silicate Naxos non-core 13.9 ± 0.3 (7)17.3 ± 0.9 (2)22.1 ± 0.8 (6)33.7 ± 0.7 (17)41.4 ± 0.7 (13)47.7 ± 0.4 (16)55.1 ± 0.8 (4)63.4 ± 0.5 (6)

NX94120 calc-silcate Naxos non-core 30.9 ± 1.9 (4)NX94106 pelite Naxos core 18.3 ± 0.2 (7)

20.6 ± 0.5 (4)NX9315 leucogneiss Naxos core 17.7 ± 0.2 (6)

28.5 ± 0.5 (4)NX9319 leucogneiss Naxos core 17.9 ± 0.2 (17)

40.7 ± 0.6 (4)NX9320 leucogneiss Naxos core 16.4 ± 0.5 (5)NX9638 migmatite Naxos core 15.2 ± 0.4 (3)

53.5 ± 1.0 (4)NX9637 pegmatitic pod Naxos core 17.4 ± 0.3 (12)SIF9345 calc-silicate Sifnos Series 43.7 ± 0.9 (5)

52.8 ± 0.8 (7)90346 qtz-phengite

schistIos Series 60.3 ± 0.7 (8)

IO9615 blueschist? Ios Series 39.2 ± 1.9 (5) 56.9 ± 1.1 (5)

6.6 SHRIMP Th-Pb dating of Monazite

Metamorphic monazite is identified only in samples from the Naxos core that have

undergone partial melting during M2b. No monazite is recognised outside the core of

Naxos or from any other island sampled. The monazite usually occurs as relatively large

grains overprinting pre-existing assemblages. They are characterised by exceedingly high

Th and U contents ranging from 65 000-180 000ppm Th and 10 000-24 000ppm U.

There is no evidence of detrital monazite in any of the sedimentary samples from the

Cyclades. Magmatic monazite is developed in Miocene S-type granitoids which intrude

Naxos (as described in Chapter 7), presumably being late-stage products of the partial

melting process.

208Pb/232Th rather than 206Pb/238U ages are quoted for monazite analyses as these are

not subject to the effects of isotopic disequilibrium, as discussed later in this Chapter and

in Appendix D.

176 Tertiary

A leucosome sample from the migmatite located in the core of Naxos (NX9637)

(Figure 6-4), was analysed and is thought to be representative of the time of partial

melting as it does not cross-cut any of the migmatite structures. The timing of

migmatisation is constrained by new zircon growth in this sample at 17.4 ± 0.3 Ma (Table

6-1). Eight of the monazite analyses define a peak at 12.9 ± 0.2 (n=8) ), which is taken

to represent the timing of monazite growth as the monazite must have formed below its Tc

(~700 ˚C). None of the age populations identified are within error of the zircon age at the

95% confidence level. The monazites from this sample have very high Th and U contents

of 130 000 ppm Th and 23 000 ppm U, but there is no correlation between Th or U

content and age, so that Pb loss associated with radiation-induced damage to the monazite

structure is not considered to be present.

a

0

1

2

3

4

10 15

NX9637n = 12

No.

of

Ana

lyse

s

Age (Ma)

4

5

6

7

10 12 14 16 18

NX94103n = 12

Age (Ma)

a) b)

Figure 6-15: Combined histograms with 1 Ma bin widths and kerned probability density curves forsamples a) NX9637 and b) NX94103.

Another migmatite sample from the leucogneiss core of Naxos, NX94103, whose

zircon protolith ages are discussed in Chapter 3, also contained monazite. Analysis of

twelve monazite grains yielded two 208Pb/232Th age populations at 12.7 ± 0.1 (n=6) and

14.7 ± 0.2 Ma (n=6) (Figure 6-15, Table 6-2). The ages are not thought to represent the

timing of partial melting that is well-constrained from the previous zircon age estimates.

No relationship between age and Th or U content in the monazites, or age and

morphology/grain size has been observed suggesting that Pb loss by fast pathway

diffusion along radiation damage structures was not a significant process in these

samples. The alternative possibilities are that the two age populations reflect two periods

Chapter 6 177

of monazite growth (Fitzsimons, 1996) or preservation of inherited radiogenic Pb in

monazite grains (Parrish, 1989). Preservation of inherited monazite ages is considered

unlikely because there is no evidence of detrital monazite in any rocks from within the

Basement or Series rocks of the Cyclades, although an earlier period of metamorphic

monazite growth preserved as inherited grains is feasible. The favoured interpretation is

that the two ages reflect separate periods of metamorphic monazite growth related to late

stage fluid movement in the Naxos core. Arguments supporting this interpretation are

detailed in Section 6.6.1.

Two samples from the wispy leucogneiss described by Buick (1988) within the

leucogneiss core of Naxos, NX9315, NX9320 contained monazite which yielded slightly

different age populations (Figure 6-16).

12 14 16

NX9315n = 11

No.

of

Ana

lyse

s

Age (Ma)

0

1

2

3

4

5

6

7

10 12 14 16 18

NX9320n = 12

Age (Ma)

a) b)

Figure 6-16: Combined histograms with 1 Ma bin widths and kerned probability density curves forsamples a) NX9315 and b) NX9320.

Zircons from both samples were also dated (Chapter 3), with NX9315 containing

new zircon growth rims at 17.7 ± 0.2 Ma (n = 6) and NX9320 containing zircons with

new growth rims dated at 16.4 ± 0.5 Ma (n = 5) (Table 6-1). Monazite from NX9315

yielded 208Pb/232Th ages of 13.3 ± 0.1 (n=9) with two older ages recognised at ~ 15 Ma

(Table 6-2). The main age population is taken as the best representation of the

metamorphic monazite age, but the two older ages are hard to explain other than by an

earlier period of monazite growth as found in sample NX94103.

Monazites from NX9320 also gave a complicated spectrum of 208Pb/232Th ages with

the main population at 14.2 ± 0.3 Ma (n = 5), with younger ages at 12.6 ± 0.3 (n = 4)

and three older grains at ca . 16 Ma (Figure 6-16, Table 6-2). The age of metamorphic

monazite development is interpreted as ca . 14 Ma from the main population of grains.

The significance of the three older ages and the smaller group at ca . 13 Ma may be

178 Tertiary

explained as the result of different periods of monazite growth. It is unlikely that

monazite growth occurred over an extended period of time as the monazites are generally

unzoned and there is no relationship between the age of the monazite and the position of

the ion probe pit.

Sample NX9438 is a folded pegmatite

which cross-cuts the Mesozoic Series rocks of

Naxos but appears to have undergone the same

deformational history. The age of this sample

can be used to constrain the maximum age of

Alpine deformation on Naxos. There is some

scatter in the ages produced but the main208Pb/232Th age population defined by twelve of

the fifteen analyses occurs at 16.9 ± 0.1 Ma,

with three younger ages forming a population at

14.8 ± 0.2 Ma (Table 6-2). As for the other

monazite samples, the two populations are both

interpreted to result from new mineral growth

related to fluid infiltration post-peak M2b.

Table 6-2: Summary U-Pb Monazite ages for Metamorphic Samples

Sample No.Spots

No.Zircons

Main Ages(No. Of Analyses)

Age Range

NX9637 12 12 12.9 ± 0.2 (8) 9.0 - 14.8NX94103 12 12 12.7 ± 0.1 (6)

14.7 ± 0.2 (6)12.1 - 15.7

NX9315 11 11 13.3 ± 0.1 (9) 12.8 - 15.4NX9320 12 11 12.6 ± 0.3 (4)

14.2 ± 0.3 (5)12.1 - 16.2

NX9438 15 15 14.8 ± 0.2 (3)16.9 ± 0.1 (12)

14.6 - 17.9

6.6.1 Comparison of Monazite ages

All the monazite ages described in the previous section fall in the range 9 - 18 Ma

with main peaks at 13-14 Ma, except for sample NX9438, a late stage pegmatite in which

the monazite formed at approximately 17 Ma. All samples show a wide range in ages

which is inconsistent with their derivation from a single age population. The possibility

of monazite inheritance from the orthogneiss basement of Naxos must be considered

(Parrish, 1990; Friedl and Finger, 1996). No monazite is found in the orthogneisses

from any of the Cycladic basement sampled in this study. Furthermore, the Cycladic

0

2

4

6

8

10

No.

of

Ana

lyse

s

15 20

NX9438n = 15

Age (Ma)

Figure 6-17: Combined histograms with 1Ma bin widths and kerned probability densitycurves for samples NX9438.

Chapter 6 179

orthogneisses are often metaluminous while monazite commonly occurs in Ca-poor

peraluminous granitoids (Cuney and Friedrich, 1987). Monazite is also not recognised as

detrital grains in any of the sediments that comprise either the Basement or the Series

rocks of the Cyclades, and so the possibility of much older monazite inheritance in

metamorphic grains is considered extremely unlikely. While the effects of small degrees

of Pb loss and/or inheritance of radiogenic Pb to explain the range of ages cannot be ruled

out, the remaining possibility is that these ages are reflecting separate episodes of

monazite growth.

The characteristically high Th and U contents found in the Naxos monazites have

occasionally been reported from late-stage magmatic products such as aplites and

pegmatites (Mannucci et al., 1986; Montel, 1993; Wark, 1993), although one instance of

“metamorphic” high-Th monazite is reported from the granulite-facies rocks of east

Antarctica (Watt, 1995; Fitzsimons et al., 1997). In this latter, a distinct 15-20 Ma age

gap between growth of melt-associated low-Th monazite and overgrowths of high-Th

fluid-precipitated monazite has been identified (Fitzsimons et al., 1997), for which the

combined operation of magmatic crystallisation and residual fluid movement is considered

to be responsible. The same processes could explain the slightly heterogeneous nature of

monazite ages found in the Naxos core, even though no core-rim structures have been

identified in the Naxos monazites. The textural setting of the monazite grains favours an

origin related to fluid infiltration, with most occurring as discrete grains close to the

boundaries of pre-existing minerals.

The fact that all monazites have very high Th contents, previously only reported

from monazites which are associated with residual magmatic fluids, suggests that the

Naxos monazites were also precipitated from residual fluids. The latter are sometimes

evident as relatively early-formed pegmatites such as sample NX9438. A fluid

precipitation origin, rather than an origin from mineral breakdown reactions, is supported

by the lack of evidence of other high REE-bearing minerals in close proximity to the

monazites. While there is no evidence of the in situ breakdown of pre-existing mineral

phases to produce metamorphic monazite, as described by Pan (1997), other mineral

phases must have reacted to release the high levels of Th and REE required to produce

new cheralite-rich monazite growth. The source of the Th and REE is not visible on a

thin-section scale so an external origin through fluid infiltration is favoured.

6.7 SHRIMP U-Pb dating of Titanite

While titanite is evident in samples from both Sifnos and Ios as an integral part of

the metamorphic assemblages, preliminary analyses indicate that nearly all of these grains

(along with many from Naxos) contain insufficient concentrations of uranium for

adequate age determinations (< 100 ppm U). However titanites from two calc-silicate

samples and one amphibolite from Naxos had sufficiently high U content to proceed.

180 Tertiary

Sample NX94121 is a calc-silicate from a high M2 grade section of the Naxos Mesozoic

Series rocks (Figure 6-4) whose sedimentary protolith is Cretaceous or younger

according to the results discussed in Chapter 4. Twenty-two analyses of eighteen titanite

grains separated from this sample yielded a range of Miocene ages (Figure 6-18). All

titanite analyses required large corrections for common Pb (70-90% - see Appendix E)

and so the selection of common Pb composition was critical and was determined

according to the method described in Appendix D. A projection of the analyses onto

Concordia results in an age of 14.5 ± 3.4 Ma (n = 20) but the large uncertainty on the

individual titanite ages does not allow separate populations to be identified.

0

1

2

3

4

5

6

7

8

5 10 15 20 25

NX94121n = 22

0.00.10.2

0.40.50.60.70.8

0.3

0 100 200 300 400 500 600 700

15 10 Ma

No.

of

Ana

lyse

s

Age (Ma)

238 U / 206 Pb

207

Pb /

206

Pb

Figure 6-18: Combined histogram with 1 Ma bin widths and kerned probability density curve fortitanites from sample NX94121. Inset is a Tera-Wasserburg concordia diagram showing all analyses.Note: the uncertainties used for the individual ages for the probability density curve shown are erronouslylow.

The ca . 14 Ma age is also recorded in new zircon growths in the same sample (see

comparisons in the ages later in this chapter). This age is thus taken as the best estimate

of titanite growth. Textures in the rock suggest that titanite might have formed at different

times in the metamorphic history of the rock, with large blocky titanites in apparent

equilibrium with the main metamorphic mineral assemblages, while small clusters of

titanite have developed within a post-peak M2b foliation related to extensional exhumation.

Only large titanite grains were analysed in this study.

Another calc-silicate from the high M2 grade Mesozoic Series rocks of Naxos,

NX94120, whose zircon ages are reported in Chapter 5, contained titanite that required

large corrections for common Pb contents (75-95% - see Appendix E). These gave an

age of 13.8 ± 2.0 Ma (n = 13) with a large MSWD of 27 using Ludwig’s calculations

(Figure 6-19, Table 6-3), but the large individual uncertainties did not allow detection of

Chapter 6 181

any age differences. No zircons of this age were found in the sample although only a

small number of analyses were made (13).

0

1

2

3

4

5

6

8 10 12 14 16 18 20 22

NX94120n = 13

0.00.10.20.30.40.50.60.70.8

0 100 200 300 400 500 600 700

10Ma15

No.

of

Ana

lyse

s

Age (Ma)

238 U / 206 Pb

207

Pb /

206

Pb

Figure 6-19: Combined histogram with 1 Ma bin widths and kerned probability density curve fortitanites from sample NX94120. Inset is a Tera-Wasserburg concordia diagram showing all analyses.Note: the uncertainties used for the individual ages for the probability density curve shown are erroneouslylow.

0

2

4

6

No.

of

Ana

lyse

s

5 10 15 20 25Age (Ma)

NX9435n = 25

0

0.1

0.2

0.3

0.4

0.5

0.6

0.7

0.8

0 100 200 300 400 500 600

10 Ma15

207

Pb /

206

Pb

238 U / 206 Pb

8

10

Figure 6-20: Combined histogram with 1 Ma bin widths and kerned probability density curve fortitanites from sample NX9435. Inset is a Tera-Wasserburg concordia diagram showing all analyses.

Titanites from an amphibolite, NX9435, from a high M2 grade portion of the Naxos

Series rocks yielded an age of 12.8 ± 2.6 Ma (n = 26) (Figure 6-20, Table 6-3). These

182 Tertiary

results required a range in common Pb corrections from 45-90% and produced a well-

defined linear array that could be extended to intersect the common Pb composition of the

sample (Figure 6-20).

Table 6-3: Summary U-Pb Titanite ages for Metamorphic Rocks

Sample No.Spots

No.Grains

Main Ages(No. Analyses)

Age Range

NX94121 22 18 14.5 ± 3.4 (20) 9.8 - 17.4NX94120 13 13 13.8 ± 2.0 (13) 11.8 - 15.0NX9435 26 24 12.8 ± 2.6 (26) 8.5 - 15.3

6.7.1 Comparison of Titanite ages

The large errors on the titanite age populations makes it impossible to compare the

titanite populations except to comment that the ages generally reflect mid-Miocene titanite

growth. The problem of high common Pb corrections inherent in dating titanite is

exacerbated in Cycladic samples by their relatively young age. This results in very low

radiogenic Pb levels and hence a large error on age determinations (Appendix D). As

mentioned in Chapter 1, most of the titanite samples analysed (30 in total) contained U

levels too low for dating purposes (< 10 ppm), but the three samples described in this

chapter are exceptional, with U contents ranging from approximately 250 - 3000 ppm.

The low levels of uranium found in most samples seems to be a problem unique to

metamorphically-grown titanite. No problems with low U content were found in the

magmatic titanite described in Chapter 7, although only two samples were analysed. The

low U contents could be due to a lack of uranium available for incorporation into the

titanites in the metamorphic environment. This would have to reflect either a low initial U

abundance in the rock or preferential incorporation of uranium into other minerals as the

abundance of uranium in rocks seems unaffected by metamorphism at grades below

granulite facies (Dostal and Capedri, 1978). A large proportion of a rock’s uranium

content may be locked into the zircon structure. Bingen et al. (1996) report that ~80% of

the uranium in an augen gneiss sample is held in the zircon. Another possibility is that U

may be excluded from the titanite structure during metamorphism. Calculations suggest

that the Si-O, Ti-O and Ca-O bond lengths in titanite decrease at high pressures (Dempsey

and Strens, 1976) and this has been confirmed in experimental studies (Kunz et al.,

1996). Compression of the CaO7 polyhedron could make substitution of high valence

cations such as U4+ or Th4+ into the Ca site more difficult.

6.8 Corrections for Isotope Disequilibrium

As Th/Pb ages are used for monazite, no adjustments to age determinations to

account for isotope disequilibrium are necessary, ignoring the potential effects of

protactinium disequilibrium (Parrish, 1990) (see Appendix D). Large adjustments to

young monazite U-Pb ages are usually required due to the production of excess 206Pb

Chapter 6 183

from the decay of high proportions of 230Th in the monazite structure. The procedure for

making isotope disequilibrium corrections to 206Pb/238U ratios measured for zircon and

titanite is described in Appendix D. In all cases here, the correction required is less than

the analytical uncertainty associated with results. A correction factor is determined by

comparing the Th/U of the host material with that of the mineral analysed to determine the

degree of initial disequilibrium. This procedure is of questionable usefulness for

metamorphic rocks where whole rock Th/U ratios do not reflect the Th/U available for

new mineral growth; the approach is really only suitable for calculating correction factors

for magmas that are pure melts (i.e. those which do not contain any restite and so can

freely exchange Th and U with crystallising minerals).

Despite these problems, correction factors are calculated here as it is the only

approach available to assess the possible effects of isotope disequilibrium in these

samples. Another approach would be to compare 206Pb/238U and 208Pb/232Th ages from

the same samples, but 208Pb/232Th ages could not be accurately calculated for either titanite

or zircon because of the extremely low Th contents found in these minerals (Appendix E).

The bulk Th/U content of two samples, one from which titanite was dated, the other from

which new zircon growth was dated, were determined by X-ray flourescence at the

Geology Department of the Australian National University and the results are listed in

Table 6-4. The Th/U in individual minerals is listed in Appendix E and an average for

each sample was taken to give an approximation of the correction factor which would be

required to account for isotope disequilibrium. In the case of the migmatitic melt pod

NX9638, only the Th/U ratios of the youngest rims of metamorphic zircon growth were

averaged as these are the ones directly related to the formation of the melt pod.

Table 6-4: XRF analyses

Sample Rb (ppm) Pb (ppm) Th (ppm) U (ppm) Reference

NX9435 23 6 6 <1 this study

NX9638 79 34 2 1 this study

Although having low levels of Th and U, XRF analyses of the titanite sample

(NX9435), indicate a bulk Th/U of ~ 6, while the average Th/U of the titanites is ~ 0.64.

Using the equations described in Appendix D, this would require a small adjustment of

only + 1.505 * 10-5 to the measured 206Pb/238U ratios to account for a deficit of 206Pb

resulting from the low initial 230Th incorporation into the titanite. This correction factor is

insignificant compared to the error on the 206Pb/238U ratios which is on the order of 10-2,

and has been ignored.

Zircon, as for titanite, can suffer from a deficit of 206Pb due to low levels of 230Th

incorporation into the mineral structure. This can potentially be a problem in metamorphic

zircons that have low Th/U values. XRF analysis of sample NX9638, one of the

migmatite samples comprised dominantly of leucosome, indicates extremely low Th and

184 Tertiary

U contents giving a Th/U ratio of ~ 2, while the average Th/U from the metamorphic-aged

zircon growths is ~ 0.001. This would require an adjustment of + 0.00048 to the206Pb/238U ratios which is considerably less than the error on these ratio determinations

(~10-3) and so no adjustment has been applied to zircons for isotope disequilibrium.

6.9 Stable Isotope Results

To examine whether some of the

metamorphic minerals dated during this

study could be related to fluid infiltration,

the stable isotope composition of eleven

samples from the island of Naxos and one

sample from Sifnos have been investigated.

Sample locations for Naxos are illustrated in

Figure 6-21B (the sample from Sifnos,

SIF9345 is described and located in Chapter

4).

Oxygen and carbon isotopes from

Naxos calcites were measured on a Finnigan

MAT 252 mass spectrometer at the

Department of Earth Sciences, Monash

University by Ian Buick. Calcite was drilled

out of handspecimens and reacted with

100% phosphoric acid in sealed vessels at

25 ˚C for two hours to liberate CO2

according to the method of McCrea (1950). Appropriate corrections have been made for

oxygen fractionation between acid and CO2 at relevant temperatures, and results are

expressed relative to V-PDB (C) and V-SMOW (O) standards. Calcite contents were

estimated from measured yields and are accurate to within ± 10%. The results are listed

in Table 6-1 and presented graphically in Figure 6-22.

Five of the eleven samples have a δ18O value greatly lowered from sedimentary

values which are expected to range from 22-25 (Baker and Matthews, 1995). This

suggests the rocks have been infiltrated by a fluid with a relatively low δ18O signature. A

range of δ13C values is found with no consistent correlation with δ18O values (Figure 6-

22). The occasionally low δ13C values found in samples with typically sedimentary

δ18O values suggests that the carbon isotopes have been influenced by decarbonation or

partial equilibration between calcite and a low δ13C reservoir such as graphite, which is

ubiquitous in pelites and also commonly found in many calc-silicate rocks.

0 1 2 km

N

Granodiorite

Gneiss/Migmatite

Upper Unit

Ultramafics

SchistMarble NX9435

NX9464

NX9461NX9463

NX94121NX86-1

86/140

85423 & 8542585-112

Figure 6-21: Location of Stable Isotopesamples from Naxos

Chapter 6 185

Table 6-1: Stable isotope composition of calcites from Naxos and Sifnos.

Sample Rock-type Wt% Calcite δ13C δ18OSIF9345 Calc-silicate† 18 -4.8 22.8NX9435 Amphibolite 8 -6.2 25.2NX9464 Calc-silicate 9 1.9 15.4NX9461 Calc-silicate 30 -2.8 23.4NX9463 Calc-silicate† 7 -5.1 17.7NX94121 Calc-silicate† 6 -3.3 14.0NX86-1* Calc-silicate 25 -2.9 16.886/140* Calc-silicate 23 -4.9 17.285425* Cal-silicate 5 1.6 22.485423* Calc-silicate 19 0.0 23.485-112* Calc-silicate 18 2.0 25.2* samples from Buick (1988), it is unknown if these display new zircon growth† samples dated during this study which show new young metamorphic zircon growth

302826242220181614121086420-8

-6

-4

-2

0

2

4

δ18O (V-SMOW)

sedimentary values

sedimentary values with low δ13C

δ13 C

(V

-PD

B)

Figure 6-22: δ13C and δ18O values from samples from Naxos and Sifnos listed in Table 6-1.

6.10 Metamorphic Fluid Composition

The development of minerals such as garnet, epidote, vesuvianite and wollastonite

in calcite-rich rocks during metamorphism is indicative of infiltration of water-rich fluids

(Tracey and Frost, 1991). The composition of fluids involved in metamorphic reactions

can be constrained by looking at the chemistry of minerals developed during fluid flow

and the composition of the fluid inclusions which they contain. To constrain the

composition ot fluids that have affected one of the samples analysed in this chapter,

NX94121, electron probe analyses were conducted using the Cameca Microbeam electron

186 Tertiary

microprobe at the Research School of Earth Sciences, Australian National University.

NX94121 is a quartz-rich calc-silicate with the post-peak M2b assemblage quartz-

clinopyroxene-plagioclase-scapolite-calcite-titanite-garnet-epidote. Calc-silicate rocks

with this mineral assemblage were interpreted by (Buick and Holland, 1991) to have ).

NX94121 has a mineral assemblage unsuitable for making P-T estimates but suitable for

calculating fluid compositions given the previous P-T estimates for these rocks by Buick

and Holland (1991). The XCO2 of the most recent fluid phase affecting the sample was

calculated using the reaction:

anorthite + calcite ⇔ grossular garnet + CO2

Fluid composition was constrained using this endmember reaction displaced by real

mineral compositions as shown in . Garnet formation was assumed to occur at 5 kbar

over a likely temp range of 500 to 600 ˚C. Mineral activities were calculated using

Program AX96 (T.J.B. Holland) using the following formulations:

calcite = assumed to be pure

garnet: 2-site mixing + activity coefficients of Newton and Haselton (1981)

plagioclase: (Holland and Powell, 1992): Darken's quadratic Formulism model 1.

The position of the reaction was drawn using the computer program

THERMOCALC (v. 2.3) (Holland and Powell, 1990) and this consistently yielded values

of XCO2 < 0.05, suggesting that the fluid was dominantly H2O. This finding is in

agreement with previous studies of fluid composition on Naxos as detailed in Section

6.3.1.

Chapter 6 187

Table 6-2: Electron probe data for garnet, scapolite and feldspar (sample NX94121).

Garnet1.1 1.2 1.3 1.4 1.5 1.6 1.7 1.8 1.9 1.10

SiO2 37.48 36.84 37.13 37.28 37.30 37.47 37.30 37.21 37.33 37.36TiO2 <0.07 0.12 <0.07 0.12 0.13 <0.07 <0.07 0.08 0.12 0.13Al2O3 21.00 20.37 20.52 20.71 21.03 20.67 20.83 20.77 20.87 21.04Cr2O3 <0.08 0.52 <0.08 <0.08 <0.08 <0.08 <0.08 <0.08 <0.08 <0.08Fe2O3 2.84 2.33 3.60 3.15 2.98 3.13 2.91 3.04 2.61 2.75FeO 19.73 17.83 18.44 18.83 19.92 18.46 19.40 19.83 19.07 19.80MnO 1.65 2.33 2.43 1.92 1.64 2.16 2.07 2.01 2.03 1.89MgO 1.45 0.88 0.61 0.81 1.59 0.56 0.86 0.69 0.71 0.80CaO 15.74 16.92 16.97 16.97 15.28 17.56 16.32 16.19 16.88 16.33Na2O <0.13 0.13 <0.13 <0.13 <0.13 <0.13 <0.13 <0.13 <0.13 <0.13K2O <0.04 <0.04 <0.04 <0.04 <0.04 <0.04 <0.04 <0.04 <0.04 <0.04Total 100.21 99.69 100.02 100.04 100.12 100.32 100.01 100.07 99.87 100.36Atomic proportions on the basis of 12 oxygensSi 2.948 2.924 2.940 2.945 2.937 2.953 2.948 2.945 2.952 2.943Ti 0.004 0.004 0.004 0.007 0.008 0.004 0.004 0.005 0.007 0.008Al 1.947 1.916 1.916 1.928 1.952 1.920 1.941 1.938 1.946 1.954Cr 0.05 0.005 0.005 0.005 0.005 0.005 0.005 0.005 0.005 0.005Fe3+ 0.168 0.214 0.214 0.187 0.176 0.185 0.173 0.181 0.155 0.163Fe2+ 1.298 1.222 1.222 1.244 1.312 1.216 1.282 1.313 1.261 1.305Mn 0.110 0.163 0.163 0.128 0.109 0.144 0.139 0.135 0.136 0.126Mg 0.170 0.072 0.072 0.095 0.187 0.066 0.101 0.081 0.084 0.094Ca 1.326 1.440 1.440 1.436 1.289 1.483 1.382 1.373 1.430 1.378Na 0.020 0.020 0.020 0.020 0.020 0.020 0.020 0.020 0.020 0.020K 0.004 0.004 0.004 0.004 0.004 0.004 0.004 0.004 0.004 0.004Total 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000

Scapolite1.1 2.1 3.1 4.1 5.1 6.1 7.1 8.1 9.1 10.1 11.1

SiO2 44.87 44.96 44.95 45.62 45.12 44.70 44.74 44.19 43.88 45.09 45.73TiO2 0.08 0.08 0.08 0.08 0.08 0.08 0.08 0.08 0.08 0.08 0.08Al2O3 29.68 29.96 29.80 29.92 29.14 29.42 29.17 29.55 29.71 30.05 36.92Cr2O3 0.08 0.08 0.08 0.08 0.08 0.08 0.08 0.08 0.08 0.08 0.08FeO 0.17 0.15 0.11 0.09 0.12 0.12 0.09 0.09 0.12 0.21 0.10MnO 0.08 0.08 0.08 0.08 0.08 0.08 0.08 0.08 0.08 0.08 0.08MgO 0.08 0.08 0.08 0.09 0.09 0.08 0.08 0.09 0.09 0.09 0.09CaO 19.02 18.95 18.82 18.97 18.17 18.98 18.65 18.73 18.64 19.07 18.40Na2O 2.57 2.68 2.62 2.70 2.70 2.53 2.62 2.60 2.55 2.70 0.56K2O 0.14 0.11 0.15 0.09 0.09 0.20 0.14 0.14 0.04 0.18 0.05Total 96.77 97.13 96.77 97.72 95.67 96.27 95.73 95.63 95.27 97.63 102.1Atomic Proportions on the basis of 16 cationsSi 6.782 6.764 6.790 6.824 6.894 6.792 6.833 6.752 6.731 6.748 6.620Ti 0.009 0.009 0.009 0.009 0.009 0.009 0.009 0.009 0.009 0.009 0.009Al 5.289 5.313 5.307 5.276 5.249 5.270 5.252 5.323 5.373 5.302 6.301Cr 0.010 0.010 0.010 0.009 0.010 0.010 0.010 0.010 0.010 0.009 0.009Fe 0.021 0.019 0.014 0.011 0.015 0.015 0.011 0.012 0.015 0.026 0.012Mn 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010Mg 0.018 0.018 0.018 0.020 0.020 0.018 0.018 0.020 0.021 0.020 0.019Ca 3.080 3.055 3.046 3.040 2.975 3.090 3.052 3.066 3.064 3.058 2.854Na 0.753 0.782 0.767 0.783 0.800 0.745 0.776 0.770 0.759 0.783 0.157K 0.027 0.021 0.029 0.017 0.018 0.039 0.027 0.027 0.008 0.034 0.009Total 16.000 16.000 16.000 16.000 16.000 16.000 16.000 16.000 16.000 16.000 16.000

188 Tertiary

Feldspar1.1 2.1 3.1

SiO2 50.48 51.03 51.15TiO2 0.08 0.08 0.08Al2O3 33.62 33.28 32.75Cr2O3 0.09 0.08 0.08FeO 0.09 0.09 0.09MnO 0.08 0.08 0.08MgO 0.09 0.09 0.09CaO 14.06 13.97 13.67Na2O 3.40 3.60 3.62K2O 0.10 0.06 0.12Total 102.1 101.5 101.7At. Prop. On basis of 8 oxygens

Si 2.251 2.285 2.286Ti 0.003 0.003 0.003Al 1.768 1.729 1.726Cr 0.003 0.003 0.003Fe 0.003 0.003 0.003Mn 0.003 0.003 0.003Mg 0.006 0.006 0.006Ca 0.672 0.656 0.655Na 0.294 0.313 0.314K 0.006 0.003 0.007Total 5.009 5.003 5.005

6.11 The Role of Fluids

Determinations of the stable isotope and chemical composition of samples from

Naxos indicate that water-rich fluids have infiltrated some rock units but not all,

presumably due to differences in permeability. This is consistent with fluid-rock ratios

found from previous studies of Tinos (Bröcker, 1990; Bröcker et al., 1993) and Naxos

(Baker et al., 1989) that suggest fluid infitration was of limited extent on these islands and

that fluid composition was controlled by the δ18O composition of the protolith lithologies

(e.g. Ganor et al., 1996). The fluids on these islands were largely internally-sourced

from dehydration of schists with low δ18O values during prograde metamorphism.

These low δ18O values are reflected in the derived fluids and cause a lowering of the

δ18O values of fluid-affected calcites (Baker and Matthews, 1994). The Naxos samples

that show the development of new zircon growth (NX94121, NX9463) are calc-silcates

that have stable isotope geochemistries consistent with fluid infiltration. However, stable

isotope data suggests that one of the calc-silicate samples which does not show new

zircon growth (NX9464) also appears to have experienced fluid infiltration. This

suggests that zircon growth, if it is associated with fluid infiltration, does not always

occur. There are few visible differences between samples NX94121, NX9463 and

Chapter 6 189

NX9464, so it is not obvious what factors might influence the precipitation of new

zircon.

The Sifnos calc-silcate sample (SIF9345) shows abundant new metamorphic zircon

growth but δ18O values are typical of sedimentary sequences. Pervasive fluid infiltration

has been identified affecting Sifnos on the basis of textural and isotopic evidence

(Matthews and Schliestedt, 1984), especially within the greenschist unit from which this

sample was taken. Matthews and Schliestedt (1984) identified the invading fluids as

being δ18O-enriched, presumably as the fluids were derived from sequences containing

metabasic rocks affected by sea-floor alteration prior to high-P metamorphism which

would give them elevated δ18O compositions (Spooner et al., 1977). Thus the fluids

associated with M2 overprinting in the greenschist unit of Sifnos are isotopically distinct

from the fluids which have infiltrated the rocks of Naxos and this is reflected in the δ18O

of the calcites. For SIF9345, the low δ13C values may give some indication that fluids

have infiltrated the sample, although the use of low δ13C as a fluid infiltration indicator is

questionable (Ganor et al., 1994).

Evidence of fluid-rock interaction in samples that also display new zircon growth is

an important consideration when interpreting the significance of ages measured from

metamorphic zircon.

6.12 Geological significance of metamorphic zircon

The meaning of metamorphic zircon ages and their relationship to P-T-t paths has

been the subject of recent debate. Marshall (1969) suggested that new zircon growth in

granulite-facies rocks was possibly associated with expulsion of a fluid phase during

falling temperatures. Roberts and Finger (1997) also suggested that zircons form during

retrograde metamorphism, and that zircon formation occurs mainly in association with

partial melting. As mentioned in Chapter 1, zirconium will often remain in solution until

the late stages of magmatic crystallisation, being concentrated in residual silicate liquids

until zircon saturation occurs (Watson, 1979; Watson and Harrison, 1983). The limited

availability of Zr means that zircon crystallisation by other means, such as breakdown

reactions or hydrothermal transport, is of very limited importance (Roberts and Finger,

1997). The relatively small body of work conducted on the conditions of zircon

recrystallisation (Pidgeon, 1992) makes it impossible to relate metamorphic zircon

produced in this fashion in any meaningful way to the P-T-t history of its host rock.

Knowing the mechanism by which zircon growth occurs is crucial for the

geological interpretation of zircon ages. There is currently little agreement concerning the

origin and significance of metamorphic zircon and the meaning of metamorphic zircon

ages (cf. Fraser et al., 1997; Roberts and Finger, 1997). Whether new zircon growth can

190 Tertiary

be related to the P-T-t history of its host rock is open to debate. As zircon has a very high

closure temperature (Lee et al., 1997), well in excess of the temperatures commonly

experienced during either metamorphism or magmatism, in most environments zircon will

record the time at which it has grown rather than the time at which it cooled through its

closure temperature. Thus the timing of new zircon growth can be directly constrained

but not the mechanism. To adaquetely interpret the significance of zircon ages, the

processes controlling new zircon growth must be distinguished.

6.13 U-Pb Ages of Metamorphic Minerals

A comparison of the U-Pb ages from zircon, monazite and titanite formed in

response to metamorphic processes on Naxos is shown in Figure 6-23. Temperatures on

Naxos are thought not to exceed 700 ˚C (Jansen and Schuiling, 1976; Buick and Holland,

1989). For the titanite-bearing samples, temperatures did not exceed 600 ± 50 ˚C (Buick

and Holland, 1991) so that all of these minerals grew approximately at or below their

closure temperatures as listed in Chapter 1. This suggests that the age of these minerals is

actually dating the time of their growth rather than the time at which they cooled below

their closure temperatures. These ages can thus be used to directly constrain the timing of

the process causing mineral growth, either by direct in situ reaction, precipitation from

melt or precipitation from fluid.

For zircons it seems clear that the youngest generation of zircons from anatectic

rocks in the core at ca . 18-17 Ma (Figure 6-23) are formed in response to the partial

melting process (as verified by the dominant age of zircons from a melt pod in the

migmatite core of the island, sample NX9637). Outside the core, zircon growth is most

likely related to sub-solidus hydrothermal processes, as no Zr-releasing breakdown

reactions are recognisable on a thin section-scale, temperatures were insufficient to

generate melting, and the morphology of the grains is inconsistent with an origin by

replacement/recrystallisation. This period of zircon growth is distinctly younger than that

found in the core, occurring at ca . 14 Ma (Figure 6-23).

Numerous stable isotope studies of the Cyclades have shown that limited

permeability and structurally-controlled fluid infiltration has occurred associated with the

M2 metamorphic overprint (Rye et al., 1976; Schuiling and Kreulen, 1979; Matthews and

Schliestedt, 1984; Baker et al., 1989; Baker and Matthews, 1994; Ganor et al., 1994;

Baker and Rutherford, 1996; Ganor et al., 1996). There is also evidence of fluid influx

during M1 in response to mineral breakdown reactions and in the protoliths prior to

metamorphism, some of which have been subjected to sea-floor alteration (Matthews and

Schliestedt, 1984). It has been suggested in previous sections that the development of

metamorphic zircon in Cycladic samples is related to fluid infiltration, with zircons

displaying complicated structures consistent with new periods of growth at distinctly

different times. In Section 6.9 it was shown that samples that contained new zircon

growth had stable isotopic signatures which were characteristic of fluid infiltration,

Chapter 6 191

consistent with the hypothesis that new zircon was hydrothermally precipitated. It is

significant that the development of new zircon growth in sedimentary rocks that have not

undergone anatexis is mainly restricted to calc-silicates. This suggests that the calc-silcates

behaved as permeable media, a suggestion supported by stable isotope evidence, but this

could also relate to the intrinsic chemical properties of the calc-silicates. Mobility of

zirconium in fluids is enhanced if the Zr can be transported as a fluoride complex (e.g.,

Rubin et al., 1993). Such solutions are mildly acidic and when transported into a reactive

medium, such as the relatively alkaline calc-silcates, may precipitate elements such as

zirconium.

If it can be shown that new zircon growth can be related to fluids, then the usual

implicit assumptions about the timing of zircon growth in relation to the closure of other

isotopic systems become invalid. This explains why some zircon ages are in close

agreement with ages derived from other systems having significantly lower closure

temperatures, without the necessity of invoking rapid cooling. This can make zircon ages

difficult to relate to P-T-t paths unless information on the fluid infiltration history of the

sample is known. Using ages derived from hydrothermally precipitated minerals, the

timing of deformational events can potentially be constrained. Structures will often focus

fluid flow, as found on Tinos (Bröcker, 1990), and dating the products of this fluid flow

places a minimum age constraint on the timing of structure development. The difference

in zircon ages between the core of Naxos, where zircon has developed in response to

anatexis, and outside the core, where zircon is hydrothermally-precipitated, could, in part,

be related to the better development of post peak-M2b extensional structures in the rocks

outside the core (Buick, 1991). These structures would act to channelise fluid flow,

possibly generated from the Naxos core during crystallisation of water-saturated partial

melts (Buick and Holland, 1991; Baker and Rutherford, 1996), and might explain the lag

in the timing of metamorphic zircon growth between the core and its surroundings.

From the arguments presented in Section 6.6.1, the development of high-Th

monazite at ca . 13 Ma is thought to occur in response to the generation of residual fluids

during crystallisation of the partial melts generated in the Naxos core. This explains the

apparent discrepancy in the youngest ages derived from zircon and monazite from the

Naxos core (Figure 6-23), as the zircon most probably grew during the partial melting

process, possibly on the retrograde P-T-t path (e.g., Roberts and Finger, 1997), while

the monazites grew approximately 4 Ma later from residual fluids, with earlier generations

of monazite at ca . 16-15 Ma possibly forming in response to long-lived fluid activity

related to residual fluid generation from partial melts in the Naxos core. Titanite ages are

relatively imprecise but suggest formation at approximately 14 Ma, in keeping with the

timing of zircon formation from the same rocks. Titanites have been generated in

response to the influx of post-M2b fluids during shearing (Buick, pers. comm.), and the

agreement of titanite and zircon ages constrains the timing of this fluid movement very

well.

192 Tertiary

0

2

4

6

8

10

12

14

5 10 15 20 25

monazite#

zircon#

titanite*

Combined Miocene Metamorphic Ages

* outside Naxos core

# within Naxos core

Age (Ma)

No.

of

Ana

lyse

s

zircon*

Figure 6-23: Combined Miocene (M2) metamorphic ages from zircon, monazite and titaniteanalysed in this study and separated on the basis of location within or outside the Naxos core.

A significant number of older metamorphic ages have been found from samples

from Naxos, Ios and Sifnos. Apart from zircon, none of the other U-bearing accessories

minerals yielded ages older than Miocene, so the following discussion of Palaeo-Eo-

Oligocene ages is specifically related to periods of new zircon growth. These periods of

new zircon development are difficult to relate to P-T conditions as they occur in response

to processes that have affected the host rock which are no longer visible, due to

overprinting by younger events. On the basis of morphologies, these overgrowths are

considered most likely to be related to periods of fluid flow. There is no evidence for

partial melting older than M2 in the Cyclades so the majority of ages must date sub-solidus

processes involving fluid transport. As found for the Miocene, zircon ages differed

between samples from within the core and from outside the core of Naxos during the

Palaeo-Eo-Oligocene. A significant period of zircon growth within the Naxos core

occurred at 28 Ma, with smaller older peaks visible in Figure 6-24, at ca . 42 Ma and ca .

54 Ma. In contrast, the most significant Palaeo-Eo-Oligocene peak for zircon growth

outside the Naxos core occurred at 48 Ma, with smaller peaks at ca . 35 Ma and 64 Ma and

in conjunction with the core zircons at ca . 42 Ma. The 42 Ma peak is also weakly evident

in zircons from the Ios Series rocks, but the dominant population in this age range is at

ca . 62 Ma. Ages from one Sifnos sample are restricted to between ca . 35 and 60 Ma,

broadly correlative with the major populations in the 50-40 Ma age range found in the

zircons from other islands

Chapter 6 193

0

1

2

3

4

5

6

7

25 30 35 40 45 50 55 60 65

within Naxos core

Ios Series

outside Naxos core

SifnosIos

Combined Palaeo-Eo-Oligocene Metamorphic Ages

Age (Ma)

No.

of

Ana

lyse

s

Figure 6-24: Combined Palaeo-Eo-Oligocene (M1) ages for zircons from Naxos, Sifnos and Ios.

The 50-40 Ma age range has been identified from other dating techniques as the

most likely timing of M1 high-P metamorphism. The significance of the other age

groupings is unclear, although it should be noted that new zircon growth can occur only

in response to external factors, be it related to fluid flow, metamorphic reactions or

recrystallisation. This suggests that significant groupings of zircon ages from different

rock units are recording the operation of tectonic processes, although it is not, as yet,

possible to precisely identify what those processes might be.

6.14 Relating Ages to P-T-t paths

The Tertiary metamorphic evolution of the Cyclades, while complicated, can be

constrained by identification of the processes involved in the generation of dateable

accessory minerals and then the combined use of information derived from different

isotopic systems. This study has focussed on U-Pb dating of minerals (zircon, monazite,

titanite) with relatively high closure temperatures. As shown in the previous section, ages

obtained from these phases can hence be interpreted as dating the timing of mineral

growth. As zircons record several different episodes of new growth which occurred in

association with previous mineral assemblages that are not preserved in most of the

samples, it is impossible to reconstruct the textural relationships of zircon growths older

than Miocene. For this reason evidence from samples that preserve evidence of pre-

Miocene metamorphism are used to help identify the relevance of zircon growth to the

metamorphic evolution of the Cyclades.

194 Tertiary

The generalised P-T path shown in Section 6.4 of this chapter is presented again

with the addition of age constraints from U-Pb dating for the main metamorphic episodes

in the Cyclades (Figure 6-25). U-Pb zircon ages presented in Chapter 5 constrain the

timing of ophiolite formation to ca . 75 Ma, which is thought to be closely followed by

associated high-T metamorphism (M?) preserved in the Cycladic Upper Unit and

constrained by argon age dating (Reinecke, 1982; Patzak et al., 1994). The timing of

hydrothermal processes related to sea-floor alteration and high-T metamorphism are

probably reflected in the ca . 78 - 68 Ma ages recorded by Cretaceous zircon overgrowths

in Series rocks of the Cyclades reported in Chapter 5. As the inferred temperatures for

M0 are higher than peak M1 temperatures, it is probable that temperatures decreased prior

to or during the elevation of pressures to eclogite-facies conditions. The closeness in

timing of ophiolite formation and high-T metamorphism, and subsequent cooling below

the argon closure temperature for white mica (~ 300 ˚C) suggests that the high-T event

was of relatively short duration, possibly associated with tectonic slicing of the oceanic

lithosphere in the early stages of the collision of Eurasia and Africa as suggested by

previous workers (e.g., Spray and Roddick, 1980). Correlating the timing of

metamorphic events in the Upper Unit, which has not experienced M1 or M2, and the

Series and Basement rocks of the Cyclades is difficult because these units have been

tectonically juxtaposed during exhumation of the high pressure rocks, so their pre-

collisional spatial relationship is difficult to constrain. For similar reasons age variations

of metamorphic minerals from different islands might reflect differences in structural level

and not necessarily record separate tectonic events. The small Paleocene zircon

populations described in this Chapter at ca . 64 and 62 Ma from the Series rocks of Naxos

and Ios, occur at the same time as rejuvenation of some argon ages in the Upper unit,

possibly due to a thermal overprint at 63-55 Ma (Altherr et al., 1979; Patzak et al., 1994).

It is unclear whether these zircon ages (which comprise a significant population in the Ios

samples) are the result of this overprint or whether they reflect the timing of devolatisation

of oceanic crust during collision and associated prograde high-P metamorphism through

the blueschist-eclogite transition. Evidence of Zr-mobility during high-P metamorphism

can be seen in the work of Philippot and Selverstone (1991) who report the existence of

Zr-bearing phases in fluid inclusions from eclogitic veins that are interpreted to represent

the composition of fluids sproduced during subduction zone metamorphism. For this

reason the 63-55 Ma zircon populations may reflect the onset of high-P metamorphism in

the Cyclades.

Chapter 6 195

ANDALUSITE

KYANITE

SILLIMANITE

P kb

ar

0

2

4

6

8

10

12

14

Temperature ˚C

zircon titanite monazite

100 200 300 400 500 600 700

13 Ma

42

54

75

? ?

M1

M2M?

D1

D2

D3

D4

15-16 Ma12

14

18

35

Figure 6-25: Schematic P-T-t path for Cycladic samples. Pathways for M1 and M2 adapted from P-T-tpaths for Series rocks of Buick and Holland (1989) , Avigad et al. (1992) , Wijbrans et al. (1993) .Pathway for Cretaceous metamorphism constrained by the P-T-t estimates of Patzak et al. (1994) for theCycladic Upper Unit.

Little evidence of peak-M1 eclogite-facies metamorphism is preserved by the

zircons. The timing of peak eclogite-facies metamorphism is constrained to be pre-54 Ma

from argon estimates (Baldwin, 1996; Baldwin and Lister, 1998) and post-75 Ma from

the time of formation of the Syros ophiolite. If the Paleocene zircon ages are the result of

M?, it could be argued that M1 was post-64 Ma. However, it is unclear whether it is valid

to make this assumption. There is a large group of zircons from both within and outside

the Naxos core and also on Sifnos formed at the same time as argon closure in fresh

eclogite samples at around ca . 54 Ma, and so this age is used to constrain the post-peak

part of the M1 eclogite-facies P-T-t path. The origin of these zircon growths is unclear as

they precede the inferred timing of eclogite-blueschist facies retrogression (thought to

occur at around ca . 54 Ma from previous work) and post-date prograde metamorphism

(as shown by the argon ages) so are probably unrelated to devolatisation of the

metamorphic pile.

The correlation of zircon ages from Naxos, Ios and Sifnos samples at 42 Ma and

their similarity to many M1 argon ages suggests this was an important period of fluid

infiltration, possibly related to fluid release from lawsonite breakdown during the down-

196 Tertiary

pressure transition from eclogite to blueschist facies metamorphism after peak M1. This is

hence used to constrain the timing of this segment of the P-T path. It is interesting to note

that Schermer et al. (1990) report evidence for the existence of two high-P events from Mt

Olympus in the Pelagonian zone at ca . 61-53 Ma and ca . 40-36 Ma from argon dating of

micas.

Continued decompression and cooling of the rocks was accompanied by

greenschist-facies retrogression, possibly in response to fluid influx from H2O released

during the breakdown of blueschist-facies minerals such as clinozoisite and paragonite

(cf. Barnicoat and Cartwright, 1997). This could account for the formation of new zircon

at 35 Ma in rocks outside the Naxos core, and at 29 Ma in rocks from within the Naxos

core. The development of greenschist-facies mineral assemblages prior to M2 would

explain the early timing (relative to peak M2) of some assemblages of this grade from

argon age constraints (Wijbrans et al., 1990). It is interesting to note however, that

insignificant new zircon growth occurred in the Ios and Sifnos samples after 42 Ma and

that these islands have not experienced the extensive M2 overprinting evident on Naxos.

A more likely possibility is that the strong population of 35-29 Ma ages on Naxos are

related to an early greenschist overprint (associated with M2?) and that the coincidence in

timing of the inferred greenschist metamorphism on Sifnos (Wijbrans et al., 1990) is

fortuitous. The position of the zircon age on the P-T diagram is hence constrained to be

somewhere in the transition between the end of M1 conditions and the beginning of M2

metamorphism.

Peak M2 conditions on Naxos were accompanied by partial melting in the Naxos

core and the formation of migmatites. Zircon is probably formed as a late stage product

of partial melting (cf. Roberts and Finger, 1997) and so the large zircon age population at

ca . 18 Ma records syn to post -peak M2b anatexis. In the core, crystallisation of melts

and/or residual fluids to the wet granite solidus over a period of approximately 5 Ma is

evident from the scattered ages of cheralite-rich monazite growth (Buick and Holland,

1991). In addition, this melt crystallisation probably supplied fluid to higher structural

levels outside the Naxos core to cause new zircon growth and development of titanite at

ca . 14 Ma. Such an interpretation would explain the apparent lag in the timing of zircon

growth between the core and the outside rocks. Fluid infiltration into these rocks may

have been aided by the development of extensional shear structures post-peak M2 (Buick,

1991). A final stage of localised fluid infiltration related to granite intrusion and contact

metamorphism (M3) is not recorded in the ages of any of the U-bearing minerals analysed

in this Chapter, but is constrained by the timing of intrusions to be discussed in Chapter

7.

Chapter 6 197

6.15 Tectonic Implications

The eclogites and blueschists of the Cyclades have been related to north-westward

subduction of oceanic and continental crust beneath the Eurasian plate (Robertson and

Dixon, 1984), during the Tertiary convergence of Africa and Eurasia (Figure 6-26).

EOCENE ~50 Ma

EURASIA

Black Sea

Pelagonian

Iran

Puturge

Bitlis

AlanyaSouth Aegean

Peloponnese

Menderes/Tauride

E. Tauride

Kirsehir

AFRICA


Sakarya

Figure 6-26: Plate reconstruction for the Eocene taken from Robertson and Dixon (1984).

As shown by the zircon ages presented in this chapter, subduction of the South

Aegean block was most likely initiated shortly after the formation of the 75 Ma Syros

ophiolite, as recorded by some of the oldest Tertiary metamorphic zircon ages at ~ 63 - 55

Ma.

6.16 Directions for Future Research

The fluid composition responsible for deposition of separate layers of hydrothermal

zircon growth found in the Cyclades is difficult to constrain. The chemistry of the fluids

associated with the latest layer of zircon growth can be constrained by looking at the

chemistry of the mineral assemblages which host it, however this does not yield any

198 Tertiary

information about the fluid compositions related to earlier episodes of zircon growth. One

way to constrain palaeo-fluid composition would be to look at the composition of fluid

inclusions contained in the different zircon growth zones. Rudnick and Williams (1987)

and Chiarenzelli and McClelland (1993) report CO2-rich fluid inclusions in zircons

formed under granulite-facies conditions which they consider representative of the

composition of an associated metamorphic fluid. This study has also identified fluid

inclusions in the outer growth rims of Cycladic zircons (which in this case are likely to be

water-rich given petrological constraints) and these have the potential to yield information

about separate fluid flow events in the Cyclades. The oxygen isotope geochemistry of

zircon (Valley et al., 1994) has recently been investigated and experimental studies of the

diffusion of oxygen (Watson and Cherniak, 1997) in zircon suggests within grain oxygen

isotope analysis could be an excellent way to constrain the thermal history of zircon

grains. Further work in these two areas would potentially be beneficial in constraining

the chemistry of multiple generations of fluids and yield important information on the

hydrothermal evolution of orogenic belts.

6.17 Synthesis

Multiple metamorphic episodes related to fluid infiltration have been identified in the

Cyclades from SHRIMP U-Pb dating of zircon overgrowths. The reproducibility of ages

from different samples from different areas of the Cyclades and their consistency with

other geological evidence suggests that zircon can be used to constrain the multi-stage

metamorphic histories of complicated orogenic belts. Zircon ages cover the entire Alpine

evolution of the Cyclades from the Cretaceous (Chapter 5) to the middle Miocene (this

chapter). As most of the Cycladic rocks have not experienced temperatures in excess of ~

600 ˚C, the formation of new zircon must be related to sub-solidus processes and this has

important implications for the interpretation of zircon ages in medium to low-grade

metamorphic terranes. Stable isotope evidence suggests that the zircon growth can be

related to hydrothermal activity and this is facilitated by the geochemical nature and the

degree of deformation of the host lithology. As the fluid infiltration history of the

Cyclades can be reasonably well-constrained, it is possible to use the time of

hydrothermal precipitation of zircon in the construction of P-T-t paths for the area. The

youngest ages recorded by new zircon growth correlate with ages from monazite and

titanite, all of which are interpreted to have formed in response to shearing related to

extension during post-peak M2b metamorphism. This study highlights the importance of

fluid interaction in the metamorphic evolution of the Cyclades.

Chapter 7 199

7. MIOCENE MAGMATIC EVOLUTION OF THE CYCLADES

7.1 Introduction

As discussed in Chapter 6, the majority of dating work on the Cyclades has been

restricted to the Tertiary history of the area with a considerable amount of data existing for

the Miocene magmatic rocks. Despite the volume of data, earlier studies have been

restricted by their reliance on mineral separates, a particular problem for conventional

TIMS U-Pb dating of zircons because of the risk of inheritance (Williams, 1992). U-Th-

Pb analyses are also complicated by the problems associated with isotopic disequilibrium

in rocks of this age (see Appendix D). To test the relationship between magmatism and

metamorphism in the Cyclades requires precise dating of the age of intrusions. To

constrain the timing of magma emplacement, preferably minerals with a high Tc should be

dated as these are more likely to record the time of melt crystallisation rather than cooling.

Regional-scale metamorphism and deformation often show a close spatial and

temporal relationship to granitoid emplacement. This was recognised by Barrow (1893)

who suggested that the Dalradian metamorphism of the Scottish Highlands was caused by

widespread contact metamorphism from mainly hidden granitoids. The idea that igneous

intrusions may control low-P high-T metamorphism is still popular (eg. Lux et al., 1986;

Oxburgh and Turcotte, 1974; Wells, 1980). An alternative explanation for the association

of metamorphism and magmatism is that they both represent responses to an increased

geothermal gradient. Such an increase may be produced by conductive heating during

crustal thickening (England and Richardson, 1977; Hollister and Crawford, 1986;

Oxburgh, 1972; Read, 1957), high mantle heat flow resulting from continental extension

(McKenzie and Bickle, 1988; Wickham and Oxburgh, 1985) or a transient deep-seated

thermal event (Hill, 1991; Sandiford et al., 1991).

The association between magmatism and metamorphism is commonly observed in

metamorphic environments ranging from low to medium pressures. The influence of

advective heat, produced by magmatic additions to the crust, in developing low pressure

(LP) high temperature (HT) Buchan-style metamorphism is well established (Barton and

Hanson, 1989; Lux et al., 1986; Wells, 1980; Wickham and Oxburgh, 1985). High

temperatures attained at relatively shallow depths are characteristic of Buchan-style

metamorphism and are inconsistent with simple thermal conduction models which fail to

predict such high temperatures (England and Richardson, 1977). The voluminous

magmatic accretion postulated to explain the origin of low-P high-T metamorphic belts

has also been applied to models for high-P high-T granulite facies metamorphism, at least

where the rocks may be shown to have undergone near-isobaric cooling (Bohlen, 1991;

Ellis, 1987; Harley, 1989; Sandiford and Powell, 1986). The link between magmatism

and moderate pressure high-T Barrovian metamorphism has been explained in terms of a

200 Miocene

magma-loading model for the Coast Plutonic Complex, British Columbia (Brown and

Walker, 1993), although most thermal models discount the influence of magmas in

generating Barrovian metamorphism (England and Thompson, 1984).

As in Chapter 6, most of the samples described in this chapter are from Naxos,

with one sample from Tinos. The concentration on samples from Naxos was primarily to

test the relationship between metamorphism and magmatism in an area displaying a range

of metamorphic grades and also with a moderately well-established chronology for the

timing of metamorphic events.


Magmatic activity in the Aegean has previously been reported as post-peak

Barrovian M2 metamorphism, with granitoid intrusion ages spanning 22-14 Ma and

associated cooling lasting until ca . 8Ma (Altherr et al., 1982). Between 22-13 Ma, huge

volumes of calc-alkaline volcanics were also erupted in the central and northeastern

Aegean (Borsi et al., 1972; Fytikas et al., 1976). However, dating the timing of

intrusions in the Cyclades has not been straightforward. It has been suggested that the

chances of precisely dating the intrusion of I-type granitoids are quite small (Schliestedt et

al., 1987) because of problems experienced when applying U-Pb systematics to these

intrusives (Henjes-Kunst et al., 1988), which meant data derived from minerals with high

Tc values were considered unreliable. Henjes-Kunst et al. (1988) interpreted all their U-

Pb ages for zircons and uranothorites from intrusives in the Aegean as the product of

complex Pb loss processes (post-igneous open system behaviour) because their ages

appeared impossibly young compared to earlier Rb-Sr, K-Ar and fission track dates on

hornblende, biotite, titanite and apatite. As a result of these relationships, it has been

suggested that young zircons can suffer a partial loss of Pb without being influenced by

any post-igneous thermal overprint (Henjes-Kunst et al., 1988). Results presented in this

chapter address and contradict this inference.

7.2.1 Naxos

Despite numerous attempts over the last fifteen years to constrain the emplacement

age of the I-type granodiorite on Naxos, workers in the area are still forced to admit that

the age is not well constrained (Wijbrans and McDougall, 1988). The post-tectonic

granodiorite on Naxos has been analysed by a number of methods. Dürr et al. (1978)

reported a Rb-Sr total rock age of 11.7 ± 0.8 Ma, as well as K-Ar and Rb-Sr (biotite)

ages of ca . 11 Ma for the granodiorite. A minimum age for the intrusion was obtained by

Andriessen et al. (1979) who determined a Rb-Sr whole rock age of 11.1 ± 0.7 Ma for

aplitic dykes cutting the main granodiorite body. The granodiorite itself does not define

an isochron (Altherr et al., 1988). A poorly defined 40Ar-39Ar apparent age plateau at

12.2 Ma (hornblende) has been taken as a close approximation of the time of granodiorite

Chapter 7 201intrusion (Wijbrans and McDougall, 1988). K-Ar biotite ages from the Naxos

granodiorite are considered to be indistinguishable from biotite ages found in the

metamorphic complex on Naxos (Wijbrans and McDougall, 1988), while hornblende K-

Ar ages from the same samples are in the range 13.6-12.1 Ma. The fact that hornblende

ages are younger in the granodiorite than in the migmatite has been used to suggest that

the granodiorite intruded post-peak M2 (Wijbrans and McDougall, 1988). A recent K-Ar

analysis of the granodiorite yielded a biotite age of 12.3 ± 0.4 Ma (Pe-Piper et al., 1997),

about 1 Ma older than previous estimates. A fission track apatite age from the main

granodiorite of 8.2 Ma (Altherr et al., 1988) is thought to record the cooling of the

intrusion to below 120 ˚C.

The only previous attempt to date the granodiorite using (conventional) U-Pb was

unsuccessful (Henjes-Kunst et al., 1988) yielding discordant ages attributed to partial Pb

loss from both zircons and uranothorites. This interpretation is strongly influenced by the

use of a two component Pb model with an inheritance of 1100 Ma assumed, and also by

the use of multi-grain samples where the choice of pristine zircon grains is not assured..

The interpretation of Pb-loss is also dependent on the presumption that zircon ages must

always be older than K-Ar or 40Ar-39Ar ages for the same sample, an hypothesis that does

not always hold true.

The timing of syn-tectonic S- and I-type granitoid emplacement, prior to the

emplacement of the Naxos granodiorite, is suggested to be 20-19 Ma from Rb-Sr whole

rock, 19.8-15 Ma from K-Ar tourmaline with younger ages of 13 Ma from Rb-Sr

muscovite, 12.7-11.4 Ma from K-Ar muscovite, and 10 Ma from apatite fission track

dating (Andriessen, 1991; Andriessen and Jansen, 1990). A recently published K-Ar age

on biotite from one of the granites in northern Naxos yielded an age of 10.1 Ma (Pe-Piper

et al., 1997).

7.2.2 Tinos

The I-type Tinos monzogranite is itself intruded by a smaller S-type granite (Altherr

et al., 1982; Henjes-Kunst et al., 1988), although Stolz et al. (1997) described this as a

microgranite. Radiometric ages from phengites in the surrounding country rock gradually

decrease with increasing proximity to the intrusion to about 19 Ma due to the thermal

influence of the Tinos monzogranite which is thought to have intruded ca . 18 Ma (Altherr

et al., 1982) on the basis of the following data. Although a Rb-Sr isochron could not be

obtained for the monzogranite, K-Ar dates on hornblende (+5% biotite) were 14.70 ±

0.25 Ma and 15.4 ± 0.6 Ma which were interpreted as minimum ages for the intrusion.

Making a correction for the biotite content in the hornblende separate increased these ages

to ca . 17 Ma, thought to reflect cooling of the intrusion. Two biotite separates gave

concordant Rb-Sr and K-Ar ages of 14.0 and 14.3 Ma. Fission track ages on titanate of

13.8 Ma and 10.8 Ma on apatite date cooling to 280 ˚C and 120 ˚C respectively. A Rb-Sr

202 Miocenewhole rock age of 14.01 ± 0.11 Ma is assumed to date the age of the small S-type granite

which intrudes the monzogranite, while hornblende from a hornfels surrounding this

intrusion yields concordant K-Ar dates of 14.7 ± 0.4 Ma and 14.8 ± 0.4 Ma (Altherr et

al., 1982). Conventional U-Pb dating on Tinos yielded ages which were considered too

young compared to other techniques and so Pb loss (resetting) was assumed to have

occurred (Henjes-Kunst et al., 1988).

7.3 SHRIMP U-Th-Pb Results

Seven samples of granitoid rocks

from Naxos, and one sample from Tinos,

were analysed to constrain the time at which

they crystallised and also the relationship

between the timing of granite magmatism

and metamorphism in the Cyclades. The

sample locations for Naxos are illustrated in

Figure 7-1.

7.3.1 Zircons

Zircons from four Naxos granitoid

samples and one Tinos granitoid were dated

in this study. The large I-type granodiorite

pluton which outcrops over much of the

west coast of Naxos was dated using both

zircon and titanite (discussed later this

chapter). NX9301 is a medium-grained to

porphyritic grey intrusive with a plagioclase-

quartz-K-feldspar-hornblende-biotite-titanite

mineralogy, and contains numerous mafic microgranular enclaves (Didier, 1973) which

reflect the interaction of the granodiorite with a more mafic magma. The granodiorite is

foliated, although this ranges from a magmatic foliation near the coast to a tectonic

foliation close to the contact with Mesozoic series rocks (Gautier et al., 1993). The

zircons in the sample are generally clear, colourless to light brown grains. They have

elongate euhedral morphologies with sharp terminations, aspect ratios greater than 3:1,

and regular oscillatory growth zoning with no evidence of older inherited cores. Twenty-

two analyses of twenty-one zircons produced a 206Pb/238U age of 12.2 ± 0.1 Ma (n = 22)

(Figure 7-2) which is interpreted to be the time at which the Naxos granodiorite

crystallised. The zircons from this sample have unusually high common Pb contents (10-

30%) but the corrections are still relatively insensitive to the initial common Pb

composition selected.

0 1 2 km

N

Granodiorite

Gneiss/Migmatite

Upper Unit

Ultramafics

SchistMarble

NX9303NX9305

NX9301

NX9439

NX9470

NX9434NX9446

Figure 7-1: Naxos Sample location map(adapted from Jansen and Schuiling, 1976).

Chapter 7 203

0

5

10

15

20

10 11 12 13 14 15 16

0

0.1

0.2

0.3

0.4

0.5

0 200 400 600 800

12 10 Ma

207

Pb /

206

Pb

238 U / 206 Pb

NX9301n = 22

Common Pb trajectory

No.

of

Ana

lyse

s

Age (Ma)

Figure 7-2: Combined histogram with 1 Ma bin widths and kerned probability density curve for sampleNX9301. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

The other three samples from Naxos are all from small intrusive bodies with S-type

characteristics (White and Chappell, 1977) although NX9303 and NX9470 are most

likely I-type intrusives which have fractionated to produce S-type characteristics.

NX9303 is a porphyritic intrusion with large phenocrysts of K-feldspar and a matrix

mineralogy of plagioclase-quartz-biotite-hornblende-titanite. The titanite was also dated

(section 7.3.3). There are two main morphological types of zircon present in the sample;

squat, slightly coloured (light brown) grains with aspect ratios less than 3:1 containing

numerous inclusions, and elongate, euhedral, clear, colourless grains with aspect ratios

up to 8:1. Neither of the zircon types appears to contain inherited cores. The elongate

grains occasionally contain hollow tubes which are thought to result from rapid

crystallisation. Twelve analyses of eleven zircons of mixed morphologies yielded

consistent 206Pb/238U ages forming two close groups at 11.3 ± 0.2 (n = 6) and 12.4 ± 0.2

(n = 6) (Figure 7-3). There is no consistent correlation between zircon age and uranium

content in the sample. A negative correlation might indicate that fast pathway diffusion of

Pb from the zircon (i.e. Pb loss) had occurred caused by radiation damage of the zircon

lattice. This type of structural damage usually only occurs in grains with high U content

(> 1000 ppm) whereas all of the granitoid samples in this study generally had much lower

U contents, and in the rare cases where a high U grain was analysed the age was not

unusually young. This indicates that Pb loss has not had a significant effect on the age

204 Miocenedeterminations. There is also no correlation between age and morphology or spot

location, so the existence of two distinct zircon popultaions is hard to explain. The two

populations may reflect variations in the time of crystallisation of zircons within the

granitoid magma, and this is the favoured interpretation of these results. As both of these

ages represent the timing of crystallisation, the younger age is taken as that best

approximating the time of granitoid emplacement.

0

1

2

3

4

5

6

9 10 11 12 13 14 15

NX9303n = 12

0

0.05

0.10

0.15

0.20

300 400 500 600 700

16 12 Ma18

238 U / 206 Pb20

7 Pb

/ 20

6 Pb

No.

of

Ana

lyse

s

Age (Ma)

Figure 7-3: Combined histogram with 1 Ma bin widths and kerned probability density curve for sampleNX9303. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

Sample NX9470 is an S-type granite with the assemblage quartz-plagioclase-

biotite-hornblende-K-feldspar. The zircons in the sample range from clear, colourless

grains to slightly discoloured inclusion-packed grains. Some of the zircons have

overgrowths that have a mottled texture full of inclusions and these are generally

concentrated at grain terminations. From twelve analyses of eleven zircons it was found

that two of these mottled zones yielded older ages ( ca . 17 and 19 Ma) than the other

analyses and had much higher common Pb contents (Appendix E), suggesting the

analyses are contaminated by Pb contained in the inclusions within these zones. The two

ages are therefore not included in the age calculation for the sample. Combining the

remaining results yields two ages of 13.3 ± 0.1 Ma (n = 4) and 15.4 ± 0.1 Ma (n = 5)

(Figure 7-4). As for the other S-type granitoid, NX9303, there is no apparent chemical

or morphological reason to explain the existence of two distinct zircon populations, and

so these are interpreted as both reflecting the time over which the granitoid crystallised.

Chapter 7 205The younger age is probably the best approximation of the emplacement age of the

sample. The oscillatory zoned inherited core of one zircon was also dated, yielding an

age of 261 Ma. This date constrains the age of the granite source to being Permian or

younger.

0

0.1

0.2

0.3

0.4

0 100 200 300 400 500 60015 Ma20

207

Pb /

206

Pb

238 U / 206 Pb

NX9470n = 12

0

1

2

3

4

5

10 12 14 16 18 20 22 24

No.

of

Ana

lyse

s

Age (Ma)

Figure 7-4: Combined histogram with 1 Ma bin widths and kerned probability density curve for sampleNX9470. Inset is a Tera-Wasserburg Concordia diagram showing all analyses. Note that one analysis at261 Ma is not included on the histogram or age probability density curve.

NX9446 is a foliated S-type granite with plagioclase-quartz-biotite and minor K-

feldspar which is thought to have intruded during the last stages of deformation on Naxos

(Buick, pers comm.) and thus places a minimum constraint on the timing of this

deformation. The zircons are of typical magmatic appearance with some broken and

resorbed cores surrounded by regular oscillatory growth zoning. Twenty analyses of

nineteen zircons revealed five anomalously old ages, three of which can be interpreted as

inherited cores at ca . 170, 300 and 315 Ma (Figure 7-5). The other three old ages at 17,

23 and 34 Ma have low Th/U but not unusually high common Pb contents and are

difficult to interpret. They may represent metamorphic zircon inherited from the

sedimentary protolith, since these ages are common for metamorphic zircon on Naxos, as

demonstrated in Chapter 6. The remaining ages form three separate populations at 12.2 ±

0.2 Ma (n = 4), 13.2 ± 0.2 (n = 7) and 14.0 ± 0.2 (n = 2). Like the previous two

granitoid samples discussed, the age of zircons from NX9446 bears no relation to their U

content, morphology or the location of the probe analysis and so Pb loss is not considered

206 Miocenean important process in generating the range in zircon ages. The four ages at ca . 12 Ma

could be related to fluid infiltration during continued shearing of the granite after

emplacement. The two older ages at ca . 14 Ma may reflect the influence of minor degrees

of mixing with an older inherited radiogenic Pb component although there is no

morphological evidence of this process. Alternatively, the ages at ca . 14 Ma may

represent zircons formed during the early stages of crystallisation of the granitoid magma.

The dominant age population at ca . 13 Ma is taken as the best representation of the time of

granite crystallisation.

0

2

4

6

8

10

12

14

0 50 100 150 200 250 300

NX9446n = 19

0.04

0.08

0.12

0.16

0.20

0 100 200 300 400 500 600 700

10 Ma152030100

207

Pb /

206

Pb

238 U / 206 PbNo.

of

Ana

lyse

s

Age (Ma)

Figure 7-5: Combined histogram with 5 Ma bin widths and kerned probability density curve forsample NX9446. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

TIN9603 is an S-type granite from

Tinos (Figure 7-6). It is porphyritic with

phenocrysts of perthitic K-feldspar in a

matrix of plagioclase-quartz-biotite-

hornblende and has allanite as a rare

accessory phase. It consists of clear,

colourless, euhedral zircons with sharp

terminations, often extremely elongate

with aspect ratios up to 10:1. Apatite

inclusions are common and many grains

contain hollow tubules which are

0 2 4 km

N

Metabasite

Granite

Marble

Schist

Upper UnitTIN9603

Figure 7-6: Tinos sample location map.

Chapter 7 207indicative of crystallisation during rapid cooling. Analysis of seventeen zircons identified

a tight cluster of ages at 14.4 ± 0.2 Ma (n = 14) with three younger analyses indicative of

later fluid infiltration and crystallisation of zircon (Figure 7-7). The age at ca . 14 Ma is

taken to represent the time of magma crystallisation.

TIN9603n = 17

0

0.05

0.10

0.15

0.20

0 100 200 300 400 500 600 700

10 Ma1520

207

Pb /

206

Pb

238 U / 206 Pb

0

2

4

6

8

10

12

10 12 14 16 18 20

Common Pb trajectory

No.

of

Ana

lyse

s

Age (Ma)

Figure 7-7: Combined histogram with 1 Ma bin widths and kerned probability density curve for sampleTIN9603. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

208 Miocene

Table 7-1: Summary of U-Pb Zircon Ages for Magmatic Samples

Sample No.spots

No.zircons


Age Range

NX9301 22 21 I-type 12.2 ± 0.1 (22) 11.6 - 13.0NX9303 12 11 I-type 11.3 ± 0.2 (6)

12.4 ± 0.2 (6)10.2 - 13.8

NX9470 14 13 S-type 13.3 ± 0.1 (4)15.4 ± 0.1 (5)

13.3 - 261

NX9446 19 19 S-type 12.2 ± 0.2 (4)13.2 ± 0.2 (7)

11.9 - 313

TIN9603 17 17 S-type 14.4 ± 0.2 (14) 12.9 - 15.8

7.3.2 Monazite

Three S-type intrusions from Naxos

have been dated using monazite Th-Pb ages.

The use of Th-Pb as opposed to U-Pb ages is

to overcome potential problems caused by

isotope disequilibrium that might affect young206Pb/238U ages. The location of the samples

is listed in Figure 7-1 and all are S-type

granites, intruded post peak-M2b (Buick,

1988). Sample NX9439 is a strongly

foliated, granitoid containing quartz-K-

feldspar- plagioclase-biotite-muscovite. Two

major 208Pb/232Th age populations can be

identified from mixture modelling at 11.6 ±

0.1 (n=10) and 12.2 ± 0.1 Ma (n = 9) (Figure

7-8). These ages are significantly different at

the 1% level and need to be explained as

separate age groups. As for the zircon ages

presented in the preceding section, the

monazite ages show no correlation between U

or Th content and age, suggesting that fast

pathway Pb diffusion in response to radiation damage of the monazite structure has not

been significant. This is despite the characteristically high Th and U content of these

monazites which are quite similar to the “metamorphic” monazites described in Chapter 6.

The same interpretation for the range found in metamorphic monazites is invoked for the

range in ages for this sample, with separate populations considered to result from an

extended crystallisation process or from the interaction of late stage fluids with the

granitoid melt. Combining the two age groups gives the preferred time of crystallisation

of the granitoid as ca . 12 Ma.

010 11 12 13

NX9439n = 19

No.

of

Ana

lyse

s

Age (Ma)

12

10

8

6

4

2

Figure 7-8: Combined histogram with 1 Mabin widths and kerned probability density curvefor sample NX9439.

Chapter 7 209

NX9305 is a well-foliated

plagioclase-quartz-biotite granite

with minor K-feldspar.

Monazites from the sample yield

a large 208Pb/232Th age group at

14.5 ± 0.1 Ma (n=9) with three

older ages forming a group at

16.5 ± 0.5 Ma (Figure 7-9).

The meaning of the older ages is

unclear, but they could reflect

radiogenic lead from a slightly

earlier phase of metamorphic

monazite growth (as described

in Chapter 6). Accordingly, the

large younger age population is

taken as the best representation

of the timing of crystallisation of the granitoid.

The third monazite-bearing S-type

granitoid, NX9434, is a relatively large,

peraluminous biotite-garnet granite with the

assemblage quartz-K-feldspar-plagioclase-biotite-

garnet-muscovite. The granitoid is foliated, but

the foliation becomes progressively weaker away

from the contacts with the country rock (Buick,

1991b). 208Pb/232Th ages from thirteen

monazites yield a well-defined population at 12.0

± 0.2 Ma (n = 11) (Figure 7-10) with two

younger ages possibly reflecting late stage fluid

infiltration. The large age population at ca . 12

Ma is taken as representative of the crystallisation

age of the S-type granitoid magma.

0

1

2

3

4

5

6

12 13 14 15 16 17 18

NX9305n = 12

Age (Ma)N

o. o

f A

naly

ses

Figure 7-9: Combined histogram with 1 Ma binwidths and kerned probability density curve forsample NX9305.

0

1

2

3

4

5

6

7

8

8 10 12 14 16

n = 13

No.

of

Ana

lyse

s

Age (Ma)

NX9434

Figure 7-10: Combined histogram with 1Ma bin widths and kerned probability densitycurve for monazites from sample NX9434.

210 Miocene

Table 7-2: Summary of U-Pb Monazite Ages for Magmatic Samples

Sample No.spots

No.zircons


Age Range

NX9439 19 11 11.6 ± 0.1 (10)12.2 ± 0.1 (9)

11.1 - 12.5

NX9305 12 12 14.5 ± 0.1 (9) 13.2 - 16.8NX9434 13 12 12.0 ± 0.2 (11) 9.2 - 12.6

7.3.3 Titanite

Titanite ages were derived from two of the I-type granite samples dated using

zircons in Section 7.3.1, which presents rock descriptions and location details for these

samples. NX9301, the I-type Naxos granodiorite, contains at least two generations of

titanite: magmatic grains and small new growths along fractures and on the margins of

pre-existing grains. The latter are thought to be related to fluid movement during the

emplacement of the granite along a progressively brittle shear zone (Buick, 1991a). All

titanite analyses require significant common Pb corrections of between 70-80 %

(Appendix E) and are sensitive to the initial common Pb composition applied, so initial Pb

is approximated using the method outlined in Appendix D. Due to the large common Pb

corrections, errors on the age derived from the intercept with the Concordia are calculated

using the method of Ludwig, as described in Appendix D. This yielded an age of 11.8 ±

0.9 Ma using thirty-eight of the analyses (Table 7-3). Variations in the titanite ages that

may be seen in the age probability density curve (Figure 7-11) do not take into account

the increase in error on individual analyses for their projection onto Concordia. This is

only incorporated in the error calculated for the titanites as a group according to the

procedure of Ludwig (1993). The significance of this age range is therefore difficult to

assess.

Chapter 7 211

0

2

4

6

8

10

12

8 10 12 14 16

NX9301

0 100 200 300 400 500 600 700

10 Ma15

238 U / 206 Pb

207

Pb /

206

Pb

1.0

0.8

0.6

0.4

0.2

0

n = 39N

o. o

f A

naly

ses

Age (Ma)

Figure 7-11: Combined histogram with 1 Ma bin widths and kerned probability density curve fortitanites from sample NX9301. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

Sample NX9303 also shows evidence of Pb loss (or U gain) from the scatter of

ages plotting to the right on the Tera-Wasserburg concordia (Figure 7-12). Analysis of

eleven grains with 60-75 % common Pb corrections produced an age of 12 ± 2 Ma

although 2 distinct peaks are recognised from mixture modelling which does not

incorporate an additional error component for the amount of common Pb correction

(Appendix D).

212 Miocene

0

1

2

3

4

5

6

9 10 11 12 13 14 15

NX9303n = 12

0

0.05

0.10

0.15

0.20

300 400 500 600 700

16 12 Ma18

238 U / 206 Pb

207

Pb /

206

Pb

No.

of

Ana

lyse

s

Age (Ma)

Figure 7-12: Combined histogram with 1 Ma bin widths and kerned probability density curve fortitanites from sample NX9303. Inset is a Tera-Wasserburg Concordia diagram showing all analyses.

Table 7-3: Summary of U-Pb Titanite Ages for Magmatic Samples

Sample No.spots

No.grains


Age Range

NX9301 39 23 11.8 ± 0.9 (38) 8.7 - 13.0NX9303 11 11 11.0 ± 2.7 (11) 11.2 - 13.4NX9303 10 10 13.6 ± 1.9 (10) 11.2 - 13.4

7.3.4 Corrections for Isotope Disequilibrium

As described in Chapter 6, no corrections for isotope disequilibrium have been

applied for monazite because only 208Pb/232Th ages have been used and these are

considered to be largely immune from the problems of disequilibrium experienced by206Pb/238U ages. For titanite and zircon, however corrections for a deficit of 206Pb might

be required and so the Th/U of samples NX9303 and TIN9603 were measured by XRF at

the Geology Department, Australian National University to determine the order of

magnitude of such corrections (Table 7-4).

Chapter 7 213Table 7-4: XRF analyses of trace element contents

Sample Rb Pb Th U Reference

NX9303 192 38 20 4 this study

TIN9603 162 34 13 6 this study

NX9301 206 41 24.7 7.3 Pe-Piper et al.

(1997)

The bulk Th/U ratio was ~ 5 for NX9303 while the average Th/U ratios for the

titanites in this sample was ~ 1.45 and for the zircon ~ 0.31. Applying the equations

described in Appendix D shows that the corrections required for both titanite and zircon in

these samples are much smaller than the error on the 206Pb/238U ratios. Titanite 206Pb/238U

ratios require an adjustment of + 1.195 * 10-5 while the error on individual 206Pb/238U

ratios is of the order of 10-3, and zircon required an adjustment of +1.6 * 10-5 while the

1σ errors on 206Pb/238U ratios are all of the order of 10-4. Similar results were found for

both titanite and zircon from the Naxos granodiorite (NX9301) using Th/U for the

granodiorite calculated from the analyses of Pe-Piper et al. (1997). As for sample

NX9303, the corrections were found to be negligble in comparison to the error associated

with the 206Pb/238U ratio. A correction of + 1.543 * 10-5 would be required for zircon

analyses whose 206Pb/238U ratios already have an error of 10-4 associated with them,

while titanite 206Pb/238U ratios would require an adjustment of +1.484 * 10-5 when the

associated error is already of the order of 10-3 (see Appendix E). Similar correction

factors are required for sample TIN9603. While Th/U ratios for the other magmatic rocks

analysed in this chapter are not available, the effects of initial disequilibrium are

considered to be negligible in comparison to the error on the 206Pb/238U ratios for these

samples.

214 Miocene7.3.5 Combined Zircon/Monazite/Titanite Intrusion Ages

0

5

10

15

20

25

30

6 8 10 12 14 16 18 20Age (Ma)

No.

of

Ana

lyse

s

Combined ResultsAll Magmatic Rocks

n = 129

Figure 7-13: Combined Miocene ages from zircon, titanite and monazite for all Naxos samples.

Combining the U-Pb ages derived from the intrusive rocks of Naxos (Figure 7-13)

clearly illustrates the restricted age range of magmatic activity on Naxos with most

intrusives aged between 13 and 11 Ma. These ages of intrusion suggest Naxos was

affected by a rapid pulse of magmatic activity during the mid-Miocene.

7.4 Discussion

7.4.1 The Effects of post-Igneous Pb loss

It has been suggested that zircons from undeformed Miocene I-type granitoids from

the Cyclades have undergone post-igneous Pb loss, which was unrelated to any thermal

overprint (Henjes-Kunst et al., 1988). Instead partial Pb loss is thought to have occurred

during rapid uplift of the Cyclades, with U gain (and possible Pb loss) associated with

deformation-enhanced fluid migrations (Henjes-Kunst et al., 1988). A few of the zircons

analysed in this chapter have undergone a minor degree of apparent Pb loss or U gain and

Pb/U ratios show a similar scatter to those reported by Henjes-Kunst et al. (1988). It

should be noted that this study has an advantage over the previous study by (Henjes-

Kunst et al., 1988) in that it utilised a much larger number of analyses made on individual

mineral grains rather than a restricted number of analyses on mineral separates, and so the

reproducibility of results is readily apparent. Due to the within-grain analysis capability

Chapter 7 215of SHRIMP which allows for careful selection of pristine zircon grain areas, and also

because of the large number of analyses possible using SHRIMP, the influence of the

relatively small number of grains affected by Pb loss or U gain is outweighed by the large

number of zircons defining a homogeneous population. It is thus possible to construct a

reasonable Concordia through the data and extract an age.

The SHRIMP U-Pb ages presented here are largely consistent with other K-Ar,40Ar-39Ar and fission track age constraints summarised in Section 7.2, as well as being

internally consistent. The agreement between the U-Pb systematics of zircon and titanite

in both samples NX9301 and NX9303 indicates that if Pb loss was important, then both

minerals must have experienced the same degree of Pb loss which seems unlikely. It is

also important to note that the U-Pb ages only appear to contradict argon ages from

hornblende, a mineral where excess argon should always be considered, and from

tourmaline, which contains only small amounts of potassium. The authors of the argon

data for the Naxos granodiorite admit that they can not accurately constrain its time of

intrusion (Wijbrans and McDougall, 1988). The apparent inconsistency between U-Pb

zircon ages and Rb-Sr whole-rock age determinations is hardly surprising. Rb-Sr whole

rock data rely on the attainment of homogeneous Sr isotope ratios, which appears to be

uncommon in initial young I-type granitoids from the Cyclades making them difficult to

date using this method (Altherr et al., 1988; Altherr et al., 1982). The same problems are

encountered in young S-type granitoids where homogenisation of Sr isotopes is

commonly incomplete as discovered by Buick (1988) and Andreissen et al. (1979) who

derived poorly constrained Rb-Sr whole-rock ages of 16 ± 8 and 30 ± 15 Ma,

respectively, for a biotite granite on Naxos.

The excellent agreement between U-Pb systematics of zircon and titanite derived

from the same samples (NX9301, NX9303) and also the more general agreement

between U-Pb systematics from zircon and titanite with Th-Pb systematics of monazite

from the granitoids of Naxos suggests that Pb loss has not influenced zircon age

determinations to a significant extent in this study, since Pb loss would be expected to

affect different phases to different extents. Data reported in this chapter are younger than

some Rb-Sr whole rock determinations, and also some K-Ar data on tourmaline and

hornblende but this is considered to be the result of problems with these dating techniques

rather than the result of the zircon U-Pb ages being impossibly young as suggested by

Henjes-Kunst et al. (1988).

7.4.2 Crystallisation/Emplacement Ages

Due to the high closure temperatures of minerals dated using U-Th-Pb systematics,

their ages provide the best approximations to the crystallisation and hence emplacement

ages of magmatic rocks. For this reason the results presented in this chapter are

considered to represent the best approximations currently available for the age of the

intrusions described.

216 Miocene7.4.2.1 The Naxos Granodiorite

SHRIMP U-Pb dating of zircon and titanite from the Naxos granodiorite yielded

ages of 12.2 ± 0.1 Ma and 11.8 ± 0.9 Ma respectively, which are within error of each

other and interpreted as good approximations for the time of granodiorite intrusion. This

is a young estimate compared to the K-Ar ages from hornblendes 13.6-12.1 Ma although

a 40Ar-39Ar age of 12.2 Ma is also obtained from hornblende and a K-Ar age on biotite

yields 11.4 Ma (Wijbrans and McDougall, 1988). The 11.1 Ma age (Rb-Sr whole rock)

of an aplitic dyke cross-cutting the grandiorite (Andriessen et al., 1979) suggests the main

intrusion had solidified by this time and cooled relatively quickly over ~ 1 Ma. Final

cooling of the granodiorite body below 120 ˚C during exhumation is recorded by a fission

track apatite age of 8.2 Ma (Altherr et al., 1982).

As discussed above, this study shows that zircon can be used to determine the time

at which Cycaldic granitoids crystallised and hence their approximate time of

emplacement, and previous arguments for discordance due to Pb-loss were based on

assumptions regarding the reliability of Rb-Sr and K-Ar ages.

7.4.2.2 Fractionated I-type granites and S-type Granites on Naxos

The other granitic bodies on Naxos all yielded U-Pb ages from zircon and titanite

and Th-Pb ages from monazite of between 12.0 - 15.4 Ma, with this range reflecting the

possible influence of fluid infiltration associated with shearing. NX9303, NX9446,

NX9439 and NX9434 all show evidence of ductile shearing but do not preserve evidence

of earlier structures recorded by their metasedimentary hosts. These late stage intrusives

yield monazite and zircon ages of 12.0 - 13.2 Ma. In sample NX9303 the age of

intrusion from zircon is verified by titanite dating as ca . 12.4 Ma. Samples NX9470 and

NX9305 are earlier intrusives, also ductilely deformed and intruded post-peak M2b at ca .

15.4 -14.5 Ma. The U-Pb ages derived from both early and late stage intrusives on

Naxos are consistent with 13 - 10 Ma ages from Rb-Sr muscovite, K-Ar muscovite and

apatite fission track dating from related samples (Andriessen, 1991; Andriessen and

Jansen, 1990). The U-Pb ages are significantly younger than the emplacement age of 20

- 15 Ma inferred from Rb-Sr whole rock and K-Ar tourmaline data for the numerous

Naxos S-type granitoids (Andriessen et al., 1979). However, as discussed above, the

application of the Rb-Sr whole rock technique on Naxos is fraught with difficulties

(Buick, 1988) and K-Ar ages on tourmaline are unreliable due to its low K. The U-Pb

ages presented in this chapter are consistent with the other K-Ar and Rb-Sr determinations

and suggest intrusion and cooling of the syntectonic granites within a period of 15 - 10

Ma.

Chapter 7 2177.4.2.3 S-type granite on Tinos

The U-Pb zircon age of 14.4 ± 0.2 Ma for the S-type intrusion on Tinos reported in

Section 7.3.1 is consistent with previous estimates of the age of this granitoid (Altherr et

al., 1982) and is considered to give the best available approximation of the intrusion age

because of zircon’s high closure temperature.

7.4.3 Relationship Between Metamorphism and Magmatism

On Naxos, most magmatism clearly post-dates M2, and foliated intrusives have

been deformed by late-stage shearing associated with the ductile-brittle transformation

during exhumation of the core complex (Buick, 1991b). The timing of peak M2 on

Naxos at ca . 18 - 15 Ma (Chapter 6) is quite distinct from the main period of magmatism

defined here at ca . 12 Ma. It is interesting to note however that the timing of intrusions

and retrograde M2 overlap with the oldest granite emplaced at ca . 15 Ma, while evidence

for fluid-related metamorphic growth of zircon and titanite can be identified at ca . 14-13

Ma (Chapter 6). This implies that magmatism accompanied the late stages of M2, with

granitoids intruding during cooling of the metamorphic complex from peak temperatures

of 700 ˚C. The additional of heat and fluids associated with the intrusion of magmas into

this level of crust could have extended the time of ductile shearing and helped to drive

retrograde reactions such as those which produced the titanite assemblages described in

sample NX94121 (Chapter 6).

The overlap between magmatism and metamorphism on Naxos is reflected in the

argon ages from the northern portion of Naxos. This area has been intruded by

voluminous I- and S-type granites yielding fairly consistent U-Pb and Th-Pb ages of ca .

12 Ma. Many of the argon ages from this area yield values in the 12-10 Ma values range

(Wijbrans and McDougall, 1988), slightly younger than argon ages from other portions

of Naxos. The apparent “younging” of argon ages towards the north of Naxos has been

interpreted as the result of differential uplift and exhumation of the metamorphic core

complex (John and Howard, 1995). In view of the time constraints now available on the

intrusion of I- and S-type granites in this region, another explanation is that the argon

ages in the metamorphic rocks have been reset by thermal perturbations associated with

these intrusions. Using the timing of intrusion of the I-type Naxos granodiorite (12.2 ±

0.1 Ma, NX9301 zircon age) which represents the most volumetrically important

intrusive, and comparing it to the last recorded timing of closure/formation of

metamorphic monazite in the core of Naxos, arguably signifying the end of peak M2

conditions (13.3 ± 0.1 Ma from NX9315), the difference in the timing of voluminous

magmatism and peak metamorphism is ~ 1.1 ± 0.3 Ma at the 99% confidence interval.

Such an age difference would be unresolvable by most dating techniques in any terrane

older than 100 Ma.

218 MioceneA direct genetic link between the Miocene intrusives currently exposed in the

Cyclades and the regional fluid-controlled retrogression that produced M2 is considered

unlikely, and stable isotope work by Brocker et al. (1990) on Tinos shows that fluids

derived from the Tinos monzogranite have a different isotopic composition from the fluids

involved in the regional greenschist facies metamorphism. Brocker et al. (1990)

suggested that an earlier phase of intrusive activity mobilized fluids and possibly supplied

heat for M2, a suggestion also made for Naxos by Wijbrans and McDougall (1988) who,

on the basis of their argon data, suggested that M2 was related to a short-lived thermal

pulse most likely associated with a deep-seated magma body.

The presence of mafic microgranular enclaves in the Naxos I-type granodiorite and

in other granitoids in the Cyclades (Altherr et al., 1982) and the relatively primitive

isotopic signatures of these intrusives (Altherr et al., 1988) is good evidence of mantle

involvement in granitoid production. The interaction of crust and mantle, and the

generation of magmas at deeper structural levels than those currently exposed, could

account for the influx of heat and externally derived fluids into the crust to produce M2.

The post-M2 intrusion of primitive enclave-bearing granitoids is evidence of melting of the

lower crust in response to an influx of mantle material. It seems logical that this process

could also generate heat and fluids that would move upwards through the structural pile

causing retrogression at progressively later times.

7.4.4 Comparison to surrounding areas

As noted by Altherr et al. (1988; 1982), the intrusion of magmas into the Cycladic

crust was accompanied by volcanic activity in the North Aegean (Borsi et al., 1972;

Fytikas et al., 1976). Young granites were also intruded post-tectonically into the

Menderes Massif of Turkey (Reischmann, 1991), suggesting that widespread magmatism

was a feature of the Aegean region during the Miocene. This magmatism has been related

to the tectonic environment of the area at this time, as discussed in the next section.

7.4.5 Tectonic Implications

In all areas of the Cyclades, granitoids have intruded post-peak M2b. The intrusives

show a consistent regional variation in modal composition, ranging from granodiorite in

the southwest (Larium, Serifos), granites in the centre (Tinos, Mykonos, Delos, Naxos)

and monzonites in the northeast (Ikaria, Kos) (Figure 7-14). This systematic trend

mainly reflects variation in the K2O content of the magmas, with K increasing with

distance from a former subduction zone located to the southwest (Figure 7-14). The

regional chemical variations seen in these granitoids is interpreted to reflect the Early

Oligocene-Miocene subduction of oceanic lithosphere beneath the region (Altherr et al.,

1982). A convergent margin setting for this area during the Miocene is consistent with

Chapter 7 219the development of high-P metamorphism at this time on Crete and in the Peloponessus

(Altherr et al., 1982).

NAXOS

TINOS

EVVIA

PELOPONESSUS

CRETE

Paleo-Subduction zone

Low/Med PMetamorphism

High P/TMetamorphism

MioceneVolcanicsGREECE

Figure 7-14: Inferred direction of subduction during the Oligo-Miocene and position of trench,highlighting the distribution of Miocene arc-related volcanics and the extent of Barrovian metamorphismin the Cyclades, and high P metamorphism developed in the External Hellenides during this time (adaptedfrom Altherr et al., 1982).

Altherr et al. (1988) have argued on the basis of field, petrological and isotopic

evidence that the granitoids in the Aegean region are the products of fractional

crystallisation of magmas produced by variable degrees of interaction between the crust

and mantle. None of the I-type grantoids sampled by Altherr et al. (1982) yielded a good

whole-rock Rb-Sr isochron, a characteristic which they interpret to be caused by

incomplete homogenisation of different sources during magma genesis. S-type granites,

which arguably display no effects of mantle involvement, are restricted to the centre of the

Aegean region on the islands of Tinos, Paros, Naxos and Ikaria corresponding to areas

with high K2O away from the former trench. In contrast to the S-types, the monzonites

found in the northeast of the Aegean display many features suggestive of a large mantle

contribution to these magmas (Altherr et al., 1988). This is inconsistent with their

tectonic position well away from the palaeo-convergent margin and has been attributed to

the development of a back-arc basin extensional regime in this area during the Miocene

220 Miocene(Altherr et al., 1988). The initiation of extension in the Aegean region at ca . 22 Ma has

been related to subduction zone roll-back of the Hellenic trench to its present-day position

south of Crete (Meulenkamp et al., 1988; Wijbrans and McDougall, 1988). This is

consistent with the southerly decrease in the age of arc-related magmatism as noted by

Wijbrans and MacDougall (1988) and ties in well with the inferred timing of M2 on Naxos

as noted by Buick (1991) and the difference in the main periods of magmatism found on

Tinos (~14 Ma) and Naxos (~12 Ma). M2 and magmatic activity in the Cyclades can both

be related to advective heat transport from mantle or lower crustal melts generated in

response to the initiation of extension in the Aegean caused by the migration of the

Hellenic subduction zone in a southerly direction.

7.5 Synthesis

The timing of magmatic activity in the Cyclades, at least on Naxos, occurs

predominantly at ca . 12 Ma as identified by mixture modelling of SHRIMP U-Pb ages

from minerals with high closure temperatures, zircon, monazite and titanite. Although the

titanite ages have relatively large errors, they show a close agreement with zircon ages

from the same samples. Monazite was analysed from different samples but also produced

ages consistent with zircon and titanite results. The consistency in ages derived from

these minerals suggests that post-igneous Pb loss did not have a significant influence on

age determinations (Henjes-Kunst et al., 1988) and the ages are in agreement with most

isotopic systems. The only discrepancies occur between the SHRIMP U-Pb age

determinations and ages from K-Ar on hornblende where excess argon could potentially

cause problems, K-Ar on tourmaline that is unreliable due to its low K content and also

with Rb-Sr whole-rock samples which have been shown to be problematic because of a

lack of homogenisation of initial isotopic ratios (Altherr et al., 1982).

The restricted range in igneous intrusion ages suggests that Naxos was affected by

a rapid pulse of magmatic activity (13 - 11 Ma) immediately following M2 metamorphism

(~ 18 Ma). While the timing of magmatism and peak-M2 metamorphism (determined in

Chapter 6) overlap within error, the main period of magmatic activity is consistently later

than the timing of peak-M2 with a resolvable difference of at least ca . 1 Ma for these

events. This small difference in ages would be difficult to resolve in terranes older than

100 Ma and so gives some indication of the close temporal relationship between

magmatism and metamorphism, indicating a rapid transition from peak metamorphic

conditions to magma intrusion. The close association between magmatism and peak-M2

metamorphism is consistent with both being the product of mantle-crust interaction

initiated by roll-back of the Hellenic subduction zone during the Miocene and the

subsequent generation of heat and fluids and the emplacement of magmas into higher

crustal levels.

Chapter 8 221

8. SYNTHESIS

The complex geological evolution of the Cyclades has been constrained by

SHRIMP U-Pb dating of accessory minerals. The integration of ages derived from three

different U, Th-bearing minerals, zircon, monazite and titanite, has provided insight into

the polymetamorphic and magmatic history of the Cyclades that has been inaccessible by

other dating techniques. As the mineral zircon can survive many cycles of crust formation

and recycling, zircon age patterns can preserve evidence of complex geological histories.

These age distributions, from the Archaean to the present, can identify the source material

from which the Cyclades are derived and can be used to reconstruct the plate tectonic

position of the Cyclades through several orogenic cycles.

A summary of all the U-Pb ages derived from this study is presented in Figure 8-1.

No.

of

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s

500

400

300

200

100

00 500 1000 1500 2000 2500 3000 3500

Age (Ma)100 200 300 400 600 700 800 900

Figure 8-1: Combined histogram with 50 Ma bin widths and kerned probability density curve for allminerals dated during the course of this study.

Mineral growth extends back to the Archaean with episodic peaks that can be related

to times of tectonic activity occurring at ca . 2900-2850, 2500-2450, 2050-2000, 1900-

1800, 1700-1650, 100-950, 900-800, 675-625, 625-525, 450-400, 350-300, 250-220

and then a range of ages down to the Miocene. The oldest rocks in the Cyclades, the

garnet-mica schists of the Basement, have a maximum depositional age of ~ 400 Ma, i.e.

Devonian. This means that any ages older than 400 Ma are inherited or detrital

components in the zircon age populations. Characterisation of these populations is

important in constraining the source areas from which the sediments and magmatic rocks

of the Cyclades were derived.

The following sections consider various stages of the geological evolution of the

Cyclades, summarising the age data and linking them with important tectonic events.

222 Synthesis

8.1 Mesoproterozoic-Archaean

The oldest protolith ages recorded in

the Cyclades date back to ca . 3200 Ma

(Archaean). There are scattered ages

throughout the Archaean and

Palaeoproterozoic, but a distinct lack of

Mesoproterozoic ages (Figure 8-2).

Comparisons with the distribution of

Precambrian ages from other areas of the

Earth’s surface (Chapter 2) reveal that this

Mesoproterozoic age gap is a diagnostic

feature of the zircon inheritance patterns of North and West African crust, and can be used

to distinguish crust derived from West Gondwana from that derived from east

Gondwana. The zircon inheritance patterns also confirm that the Cyclades most probably

formed part of the northern margin of Africa prior to the Jurassic. Correlations between

the Cyclades and other parts of the inferred North African margin, such as the Menderes

Massif and the Pelagonian zone are now possible based on this dataset.

8.2 Neoproterozoic

In the early Neoproterozoic, age

peaks are defined by small numbers of

zircons in the 1000-800 Ma age range,

corresponding to a time of extensive

volcanic arc activity in northeast Africa.

This rift-related volcanism was possibly

related to the breakup of the supercontinent

Rodinia. The first large (n > 10) age

populations are found in the 650-550 Ma

age range (Figure 8-3). Two age clusters

form at approximately 640-600 Ma and

575-550 Ma. These two age peaks are common to many of the crustal segments that

comprised the supercontinent Gondwana and are consistent with the timing of the “Pan-

African” orogeny. This orogeny marks the time at which East and West Gondwana

collided, and the two ages are widespread throughout the African continent. The

distinction between early Pan-African (640-600 Ma) and late Pan-African (575-550 Ma)

ages has been noted by other workers and may reflect different periods of collision during

the consolidation of the Gondwanan supercontinent. The younger Pan-African ages

0

2

4

6

8

10

1000 1500 2000 2500 3000 3500Age (Ma)

No.

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sFigure 8-2: Archaean-Mesoproterozoic ages.

0

2

4

6

8

10

12

14

500 600 700 800 900 1000Age (Ma)

No.

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Figure 8-3: Neoproterozoic ages.

Chapter 8 223correspond to the age of the orthogneissic basement of the Menderes Massif in western

Turkey, suggesting this was the time of extensive granite generation in the region.

8.3 Early Palaeozoic

The largest clustering of ages in the

Early Palaeozoic occurs between 450-400

Ma, corresponding to the timing of the

Caledonian orogeny in northern Europe,

although it is unrelated to this episode.

These ages are not found in the Menderes

Massif of Turkey nor in most areas of

Turkish crust suggesting an important

difference between these sections of the

Apulian-Anatolian plate during the

Palaeozoic .

8.4 Permo-Carboniferous

Differences in age distributions

between the Cyclades and the Menderes

Massif are also evident in the Permo-

Carboniferous. While the basement of the

Menders Massif is latest Proterozoic to early

Palaeozoic in age, the Cycladic basement is

dominated by 330-300 Ma orthogneisses.

The dominant 330-300 Ma peak in ages

found in the Cyclades represents a period of

extensive magmatic activity in this area and

parts of mainland Greece such as the

Pelagonian zone. Magmatism coincides with the time of collision between Gondwana

and Laurasia and associated with voluminous S-type granite intrusion throughout Europe

related to the Variscan orogeny.

0

1

2

3

4

5

6

7

350 400 450 500 550Age (Ma)

No.

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s

Figure 8-4: Early Palaeozoic ages.

0

5

10

15

20

25

30

35

240 260 280 300 320 340 360Age (Ma)

No.

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s

Figure 8-5: Permo-Carboniferous ages.

224 Synthesis8.5 Triassic-Jurassic

The Variscan orogeny was followed

by a period of extensive rifting,

sedimentation and volcanic activity in the

Triassic (240-220 Ma) related to the

opening of Tethys and recording the

earliest stages of the break-up of the

supercontinent Pangea. This tectonic

activity is reflected in the ages of

magmatic zircons found in the Cyclades

that show a large clustering of ages in the

250-220 Ma age range. Rifting of

continental blocks, such as the Cyclades, from the northern margin of Gondwana began

at this time. Triassic volcanism and associated sedimentation is widely reported from

other areas of the Hellenides and can now be confirmed for the Cycladic region.

8.6 Cretaceous

The separation of continental blocks

from the northern margin of Gondwana

continued in the Cretaceous despite overall

convergence of the African and Eurasian

plates. Numerous Cretaceous metamorphic

zircon overgrowths can be identified

suggesting that the Cyclades were

undergoing some form of active tectonism

at this time. Oceanic basins, formed in the

wake of rifted continental blocks from the

northern Gondwana margin during the

Jurassic and Cretaceous, were eventually closed in the Late Cretaceous causing obduction

of ophiolites and high-T metamorphism. A record of these events is preserved in the ages

of zircon rims and also in the Cycladic Upper Unit. The similarity in ages between the

ophiolite sequences preserved in the Upper Unit and the high-P ophiolite sequence of

Syros, dated in this study at ca . 75 Ma, suggests that the Upper Unit may represent an

excellent analogue to the Series rocks of the Cyclades. As the Upper Unit has not

experienced the Alpine high-P metamorphism that the Series rocks have undergone, it can

provide useful information on the pre-Alpine history of the Series.

0

5

10

15

20

25

140 160 180 200 220 240 260Age (Ma)

No.

of

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sFigure 8-6: Triassic-Jurassic ages.

0

2

4

6

8

10

60 80 100 120 140Age (Ma)

No.

of

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s

Figure 8-7: Cretaceous ages.

Chapter 8 2258.7 Tertiary metamorphic evolution

The Palaeo-Eocene collision of the

African and Eurasian plates resulted in high-

P metamorphism of the Cycladic Basement

and Series rocks. A record of this

complicated geological evolution is

preserved by a sequence of metamorphic

zircon rim ages formed in response to fluid

activity associated with the collisional

process (M1). These zircon rims, mainly in

the range 65-40 Ma, can be identified as

metamorphic in origin by their

morphologies and very low Th/U ratios. They are notably absent from the zircons of the

Basement, indicating that an important lithological control is governing new zircon

growth, most probably permeability. The ages found in the Series rocks record the final

closure of relicts of the ancient Tethys ocean. Overprinting of the Series rocks by

Barrovian amphibolite to greenschist facies grade metamorphism (M2) was also

accompanied by new zircon growth in the age range 35-14 Ma. Partial melting in the

Naxos core was accompanied by new zircon growth at ~ 18 Ma., while the fluid

movement associated with the final stages of ductile shearing on Naxos is recorded by the

growth of new zircon, titanite and monazite at ~ 14-13 Ma.

8.8 Miocene magmatic activity

Magmatic activity on Naxos

dominantly occurs at ca . 12 Ma, although

some small intrusions as old as 15 Ma are

also recorded. As found in most Barrovian

metamorphic terranes, the timing of

magmatism closely post-dates the timing of

peak metamorphism (M2b). The

relationship between magmatism and

metamorphism is not well understood,

although the intrusive rocks have provided

heat and fluids to drive at least localised

metamorphic reactions (M3). The agreement between zircon, monazite and titanite ages

from the Naxos intrusions suggests that Pb loss has not effected the ages derived from

these minerals, and that instead they reflect the timing of crystallisation/emplacement of

the magmatic rocks.

0

5

10

15

20

25

30

35

20 30 40 50 60 70Age (Ma)

No.

of

Ana

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s

10

Figure 8-8: Tertiary metamorphic ages fromzircon, titanite and monazite.

0

10

20

30

40

50

60

70

5 10 15 20 25Age (Ma)

No.

of

Ana

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s

Figure 8-9: Miocene magmatic ages fromzircon, monazite and titanite.

226 SynthesisFrom this myriad collection of ages, we can conclude that U, Th-bearing accessory

minerals having high blocking temperatures, particularly zircon, can yield important

information about the metamorphic and magmatic history of complicated geological

environments. Zircon is seldom used to constrain the P-T-t paths of high-P, low-T and

medium-P, medium-T metamorphic terranes but this study demonstrates that with careful

identification of metamorphic overgrowths, zircon ages can successfully be integrated

with P-T information. The relatively young age of the Cyclades makes it an ideal area to

apply geochronology to constrain the timing of tectonic processes because the normal

percentage uncertainties associated with age values amount to a relatively small amount of

time. This means that it is possible to resolve events that are separated by time differences

on the order of 1 Ma. This age resolution has enabled the identification of multiple

episodes of metamorphic zircon growth. The small volume of this new growth and its

relatively low radiogenic Pb levels (due to its young age) could be successfully dated only

using a technique that allows high spatial and depth resolution such as the SHRIMP ion

microprobe. The within-grain analysis capability of SHRIMP has also allowed the early

history of the Cyclades to be constrained by the dating of detrital and inherited zircon

components as old as Archaean. SHRIMP dating of U, Th-bearing accessory minerals

can be applied to a number of different tectonic problems such as the age of magmatism

and metamorphism, the provenance of sediments and the character of granite source

regions. This information can be particularly useful in testing the validity of past plate

reconstructions and can provide valuable information about the geological evolution of

any terrane.

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Appendix A 227

A : PUBLISHED/SUBMITTED WORKS

A 1. Submitted papers

Keay, S . Steele, D. A & Compston W. (in review) Identifying granite sources:

Evidence against a Precambrian basement to the Lachlan Fold Belt, Eastern Australia.

Submitted to Contributions to Mineralogy and Petrology.

A 2. Other Publications

Keay, S. & Lister, G. (1996) Evolution of the Naxos core complex, in Lister. G.

and Forster, M. (eds) Inside the Aegean Metamorphic Core Complexes, Australian

Crustal Research Centre Technical Publication 45: 61-74.

Keay, S . & Lister, G. (1996) Inside the dome of the Naxos core complex, in

Lister. G. and Forster, M. (eds) Inside the Aegean Metamorphic Core Complexes,

Australian Crustal Research Centre Technical Publication 45: 75-88.

Keay, S . & Lister, G. (1996) The Naxos detachment fault, in Lister. G. and

Forster, M. (eds) Inside the Aegean Metamorphic Core Complexes, Australian Crustal

Research Centre Technical Publication 45: 89-94.

Moore, L., Lister, G. & Keay, S . (1996) Thera: the core complex that became a

volcano, in Lister. G. and Forster, M. (eds) Inside the Aegean Metamorphic Core

Complexes, Australian Crustal Research Centre Technical Publication 45: 21-26.

Lister, G. & Keay, S. (1996) The lower plate of the Ios core complex, in Lister.

G. and Forster, M. (eds) Inside the Aegean Metamorphic Core Complexes, Australian


Lister, G. & Keay, S . (1996) The Paros detachment and its mylonites, in Lister.

G. and Forster, M. (eds) Inside the Aegean Metamorphic Core Complexes, Australian


228 Published/Submitted Works

A 3. Conference Abstracts:

Keay, S., Lister, G.S.L. & Compston, W. (1996) Thermal pulses and Barrovian

metamorphism, 13th Australian Geological Convention, Canberra, Geological Society of

Australia, Abstracts No. 41, p. 228.

Keay, S . , Compston, W. & Lister, G.S.L. (1995) U-Pb dating of metamorphic

minerals: Evidence for the transience of metamorphic processes? Australian Conference

on Geochronology and Isotope Geosciences 3, Curtin University of Technology, Perth,

W.A. p. 13.

Keay, S . & Steele, D.A. (1994) The chronology and origin of S-type granites

from north-east Victoria: Constraints from SHRIMP U-Pb zircon systematics, 12th

Australian Geological Convention, Perth, Geological Society of Australia, Abstracts No.

37, p. 208.

Appendix A 229

Following is a collection of my submitted and published works.

The first is a paper submitted to Contributions to Mineralogy and Petrology in

August, 1997 and currently in review. It represents work undertaken during the first six

months of my PhD candidature which focussed on the U-Pb systematics of zircons from

Lachlan Fold Belt granites. I. Williams was the original supervisor for this work

however the paper was written with my current supervisor W. Compston who assisted

with data reduction and interpretation of results. The second author D. Steele, provided

the samples and the geological background for the work from an idea we developed

jointly for testing the origin of the granites in north-east Victoria.

Also included are six chapters from “Inside the Aegean Metamorphic Core

Complexes” an Australian Crustal Research Centre Technical Publication No. 45 edited

by G. Lister and M. Forster. This is a field guide written to accompany a week-long

post-conference excursion which I co-led with G. Lister and M. Forster around the

Cyclades. The field-trip was organised to link in with the 1996 Penrose conference on

Exhumation processes held in Crete. I am first author on three chapters (9, 10 and 11)

which are substantially my own work, complemented by detailed structural observations

and interpretations of one of my supervisors, G. Lister. I am second author on two

chapters (5 and 11) with G. Lister, to which I contributed ideas and observations about

the geology of the areas described. I mainly supplied the historical background for

Chapter 3 which includes me as third author.

Sample Location and Description 231

B : Sample Location

SampleName

ANUNo .

SHRIMPmount

Description Map Reference

Samples from which zircons were analysedIO9403 97759 Z1978 Ios Orthogneiss 25° 16’E 36° 44’NIO9404 97760 Z1978 Ios Orthogneiss 25° 16’E 36° 44’N89640 89640 Z2405 Ios Orthogneiss 25° 18’E 36° 43’NIO9607 97761 Z2665 Ios Leucogneiss 25˚ 17.55’E 36˚ 42.56’NIO9606 97762 Z2665 Ios Garnet Mica Schist 25˚ 17.55’E 36˚ 42.56’NIO9609 97763 Z2665 Ios Garnet Mica Schist 25˚ 17.43’E 36˚ 42.95’NPA9606 97764 Z2644 Paros Orthogneiss 25˚ 12.71’E 37˚ 07.75’NPA9601 97765 Z2665 Paros Orthogneiss 25˚ 09.26’E 37˚ 06.10’NSK9601 97766 Z2633 Sikinos Orthogneiss 25° 08’E 36° 39’NNX9314 97767 Z1889 Naxos Layered Acid Gneiss NAXOS 7642.3F

E1180 N21125NX9485 97768 Z2645 Naxos Layered Acid Gneiss NAXOS 7642.1E

E1830 N17650NX9315 97769 Z2264 Naxos Leucogneiss NAXOS 7642.3F



E0005 N20990NX94103 97772 Z2153 Naxos Migmatite NAXOS 7642.2E

E2160 N16920NX9638 97773 Z2665 Naxos Migmatite 25° 30.64’E 37° 07.67’N NX9637 97774 Z2782 Melt Pod Naxos Migmatite 25° 30.59’E 37° 07.73’N NX9451 97775 Z2156 Naxos Quartzite NAXOS

E5400 N30940NX9481 97776 Z2217 Naxos Quartzite NAXOS 7642.1E

E2210 N17170SY9603 97777 Z2665 Syros “Vari” Orthogneiss 24° 58’E 37° 24’N89642 89642 Z2405 Syros Retrogressed Eclogite 24° 55’E 37° 30’N89646 89646 Z2405 Syros Quartzite 24° 54.5E 37° 29’NSY9630 97778 Z2644 Syros Schist 24°54’E 37° 28’NNX9461 97779 Z2298 Naxos Calc-silicate NAXOS

E6110 N32330NX9463 97780 Z2158 Naxos Calc-silicate NAXOS

E7960 N29630NX9490 97781 Z2264 Naxos Pelite NAXOS

E3800 N34000NX9464 97782 Z2038 Naxos Calc-silicate NAXOS 7642.2E

E3965 N17830NX94112 97800 Z2298 Naxos Calc-silcate NAXOS 7643.2E

E10520 N18250NX94120 97783 Z2613 Naxos Calc-silicate NAXOS 7633.3B

E4570 N10360NX94121 97784 Z2155 Naxos Calc-silicate NAXOS 7632.4B

E3050 N9860NX94106 97785 Z2298 Naxos Pelite NAXOS 7642.2E

E0250 N1721089639 89639 Z2405 Ios Glaucophane Schist 25° 15’E 36° 45’NIO9615 97786 Z2644 Ios Gt-glaucophane Schist 25˚ 16.20’E 36˚ 43.69’N90346 90346 Z2405 Ios Qtz-phengite Schist 25° 15’E 36° 45’N

232 Appendix B

SampleName

ANUNo .

SHRIMPmount

Description Map Reference

FL9602 97787 Z2633 Folegandros Pelite 24° 54’E 36° 37’NSK9603 97788 Z2633 Sikinos Metabasic Schist 25° 10’E 46° 43’NSIF9345 97789 Z2363 Sifnos Calc-silicate 24° 44’E 36° 56’NNX9301 97790 Z1870 Naxos I-type Granodiorite NAXOS 7641.4F

E6510 N20680NX9303 97791 Z2298 Fractionated S-type Granite E6740 N7700

NAXOS 7633.1ANX9470 97792 Z2613 Naxos I-type Granitoid NAXOS 7633.7D

E4530 N15275NX9446 97793 Z2613

Z2644Naxos S-type Granite NAXOS 7632.6C

E0050 N12520TIN9603 97794 Z2665 Tinos S-type Granite 25° 34’E 37° 34’NSamples from which monazite was analysed (denoted by suffix “M”)NX9637M 97774 Z2037 Naxos S-type Granite 25° 30.59’E 37° 0773’NNX94103M 97772 Z2922 Melt Pod Naxos Migmatite NAXOS 7642.2E

E2160 N16920NX9315M 97769 Z2922 Naxos Leucogneiss NAXOS 7642.3F

E0940 N21240NX9320M 97771 Z2922 Naxos Leucogneiss NAXOS 7642.3F

E0005 N20990NX9438M 97798 Z2301 Pegmatite NAXOS 7632.6C

E0510 N12460NX9439M 97795 Z2037 Naxos S-type Granite NAXOS 7632.6C

E0740 N12290NX9305M 97796 Z2301 Naxos S-type Granite NAXOS 7632.4B

E2000 N10850NX9434M 97797 Z2301 Naxos S-type Granite NAXOS 7632.6C

E2390 N12030Samples from which titanite was analysed (denoted by suffix “T”)NX94121T 97784 Z2155 Naxos Calc-silicate NAXOS 7632.4B

E3050 N9860NX94120T 97783 Z2615 Naxos Calc-silicate NAXOS 7633.3B

E4570 N10360NX9435T 97799 Z2265 Naxos Amphibolite NAXOS 7632.6C

E2210 N12180NX9301T 97790 Z1858

Z2313Naxos I-type Granodiorite NAXOS 7641.4F

E6510 N20680NX9303T 97791 Z2313 Fractionated I-type Granite E6740 N7700

NAXOS 7633.1A

Appendix C 233

C : SAMPLE PREPARATION

C 1. Rock Crushing

Australian custom regulations required all rock samples to be cleaned with

concentrated bleach using a scrubbing brush to remove any organic material on the rock

surface before being shipped to the country. Prior to mineral seperation, rock samples

were split to remove weathered material and then 1.5 kg of fresh rock passed through a

jaw crusher to form chips less than 2 cm3. The chips were washed using tap water and

then crushed into a powder using a tungsten carbide rock mill. The powder was sieved

through a 250 µm mesh and washed in tap water to remove clay-size particles and dried

under heat lamps in a closed fume cabinet.

C 2. Mineral Separation

Heavy minerals were separated from the rock powders using standard heavy liquid

separation techniques in specially designed down-draft fume cabinets. The sample was

first stirred into a funnel containing tetrabromomethane (sg 2.96 g.cm-3) to float off

minerals with low specific gravities such as quartz, feldspar and micas. Some of this

light fraction was removed and cleaned for potential analysis by 40Ar-39Ar while the rest

was discarded. The heavy fraction was poured out of the base of the funnel into filter-

paper lined funnels where it was thoroughly cleaned using acetone. The sample was then

stirred into methylene iodide (sg 3.3 gcm-3) to remove minerals such as hornblende

which float. Heavy minerals were extracted from the base of the funnel and cleaned with

acetone. Individual minerals were further separated from this heavy fraction by virtue of

their magnetic properties. Strongly magnetic heavy minerals (such as magnetite) were

removed using a hand magnet and the sample was then passed through a Frantz

Isodynamic Seperator. Monazite was separated at 1 Amp and 10˚ tilt, titanite at 2 Amps

and 2˚ tilt while essentially non-magnetic zircon was concentrated in the non-magnetic

fraction at 2 Amps and 2˚ tilt. These separates were then hand-picked under a microscope

to ensure purity.

C 3. SHRIMP Mount Preperation

SHRIMP has the advantage over conventional mass spectrometry techniques of

being able to perform in situ analysis of geological material. Little sample preparation is

required: samples must have polished surfaces and been cut to a size which will fit into a

SHRIMP sample holder (25mm diameter). Polished sections and blocks can be used for

in situ work, or alternatively individual grains can be concentrated and mounted in epoxy

234 Sample Preparation

and then polished using diamond paste until the cores of the grains are exposed.

SHRIMP mounts (either polished disc or epoxy mounts) were then cleaned thoroughly

using a biological detergent and then ethanol before being rinsed in distilled water and

dried in an oven at 50°C for over one hour. Clean and dry mounts were then handled

with rubber gloves and given a conductive coating of gold under vacuum to dissipate any

charge buildup during analysis using SHRIMP. Standards for measurement of different

minerals were either mounted in epoxy with the sample or mounted separately. (Unlike

SHRIMP I, SHRIMP II can accomodate dual mounts, and hence polished sections were

analysed with this instrument so that a standard could be measured in an adjoining

holder). Standards used include; zircon SL13 from Sri Lanka, monazite from the

Delegate Adamellite, titanite from the Khan pegmatite and rutile from Kragero, Norway

(see Appendix D for further discussion of the use of these standards).

C 4. SHRIMP Mount Imaging

Before final cleaning and coating, the mount or polished sections were

photographed in both reflected and transmitted light using an optical microscope, with

more detailed imagery being conducted using scanning electron microscopy (SEM).

Back-scattered SEM images of both titanite and monazite were obtained to reveal any

heterogeneities in the internal growth structure of the samples, while zircons were imaged

using cathodoluminescence (CL). CL images were obtained using the Australian National

University Electron Microscopy Unit’s Hitachi S2250 SEM, fitted with a curved mirror to

reveal internal structures in the zircons. The intensity of light reflected in

cathodoluminescence images has been ascribed to intrinsic zircon luminescence, REE

content, vacancy concentration and crystal lattice damage. In general the intensity of light

seems to be inversely proportional to the uranium content in the grain.

Back-scattered electron images of titanite, monazite and zircon were obtained using

a Cambridge S360 electron microscope at the ANU’s Electron Microscopy Unit. The

grey-scale in back-scattered electron images was proportional to the average atomic

number so heavy elements such as U (n = 92) which were concentrated in minerals such

as zircon will make their host mineral appear very pale in colour in comparison to

minerals containing lighter elements.

Appendix D 235

D : ANALYTICAL PROCEDURE

D1. Radioactive Decay

Naturally occurring radioactive isotopes have unstable nuclei that spontaneously

decay by emission of energy or particles to form different isotopes until a stable daughter

product is formed. The law of radioactive decay states that the number of atoms

disintegrating per unit time (dN/dt), the decay rate, is proportional to the total number of

radioactive atoms present:dN

dtN∝

We can define a constant of proportionality, the decay constant λ, for different

parent elements:dN

dtN= − λ

The decay constant is simply a measure of the probability that an atom will decay

within a set time period. We can calculate the number of daughter atoms produced over

different periods of time by rearranging and then integrating the above expression to

produce the decay equation:

- ∫ = ∫dN

Ntλ ∂

Giving − = +ln( ) N t cλWhen decay begins at t0 , the number of parent atoms is N0, such that

c = − ln(N0 )

So − ln(N) = λt− ln(N0 )

And −λt = ln(N / N0 )

Take the exponent e−λt = N / N0 to obtain the

Decay Equation : N = N0e−λt

The number of radioactive atoms present in the original sample ( N0), cannot be

measured so to calculate the age of the sample, the accumulation of stable daughter atoms

( D) is measured and an accumulation equation may be derived: N D N0 = +

N D N e t0 0 = + −λ from which we obtain the

Accumulation Equation : D N e t ( )= −λ 1

This equation can be expressed in terms of t to define the age of the sample.

e

D

Ntλ = + 1

then taking logs yields the

Age Equation : t

D

N ln( )

=+ 1

λ

236 Analytical Procedure

The age of a sample can hence be determined by measuring the concentration of

daughter and parent atoms present today (time t), and using the experimentally determined

decay constant (λ). Two assumptions are implicit when using the Age Equation:

1. When a mineral forms it has only parent and no daughter atoms, or if daughter atoms

are present these can be seperated from the total number of atoms present.

2. The system is closed, ie. no parent or daughter atoms are added or lost to the system.

D2. U-Th-Pb Geochronology

Uranium has several radioactive isotopes of which 238U (99.27%) and 235U

(0.72%) both decay to different stable isotopes of lead, 206Pb and 207Pb respectively, via

a chain of decay through relatively short-lived radioactive intermediate daughter products.

Similarly the most abundant radioactive isotope of thorium, 232Th, decays to a stable lead

isotope, 208Pb (Table D-1). Three independent ages can be assessed by measuring the

isotope ratios from these three decay schemes. As two isotopes of U, with different half-

lives, decay to form two independent isotopes of Pb, the isotopic composition of Pb can

also be used to assess the age of a sample using the 207Pb/206Pb ratio. The 207Pb/206Pb

ratio is insensitive to modern Pb loss, since it is assumed that any Pb lost has the same

isotopic composition as the measured Pb, assuming that the present 238U/235U ratio is

constant.

The 207Pb/206Pb ratio is related to the decay of the original uranium as follows

207 206 235 238235

238

1

1Pb Pb U U

e

e

t

t* *

=−

−

λ

λ where

the asterisk denotes a radiogenic daughter isotope.

Table D-1: Decay schemes for uranium and thorium isotopes.

Decay System λλλλ t1

2

238U => 206Pb + 8 α + 6β- + Q 1.55125 ×10−10 yr−1 4468 Ma235U => 207Pb + 7 α + 4β- + Q 9.8485 ×10−10 yr−1 704 Ma232Th => 208Pb + 6 α + 4β- + Q 4.9475 x 10-11 yr- -1 14.01 Ga

The different ages calculated from the measured Pb/U ratios and the 207Pb*/206Pb*

ratio are the same, then the isotopic ages are said to be “concordant”. This can be

assessed using a Concordia diagram such as the Wetherill (1956) (Figure D-1) or Tera

and Wasserburg (1972) (Figure D-2) diagrams. Analyses that plot on the curves of these

diagrams have remained closed to isotopic disturbance. “Discordant” analyses that plot

off the curves indicate that the isotopic systems have been disturbed, ie. opened to

Appendix D 237

external isotopic exchange. If the samples were disturbed by a single later event, either

type of Concordia diagram can be used to assess the original age of a sample and the time

at which the later isotopic disturbance occurred. The discordant analyses will form a

linear array termed a Discordia chord which will intersect the Concordia curve at the age

of the sample and the age of the disturbance. If the isotopic systems have been opened to

exchange more than once, or if mixed age populations are sampled, then the analyses will

not define a simple discordia line but will show complex and irregular scatter.

Concordia

Discordia

207 Pb* / 235 U

206Pb*238U

1.0

1.5

2.0

2.5

3.0

3.5

T'

Q

QT0

0 2 4 6 8 10 12 14 16 18 20 22 24 26 28 300

0.1

0.2

0.3

0.4

0.5

0.6

0.7

0.8

0.9

1.0

T

Figure D-1: Effects of episodic Pb loss (or uranium gain) on a Wetherill Concordia diagram. Amineral which has lost all Pb would plot at the origin. If Pb loss occurred at time T’ the system wouldfollow a straight line (discordia) to the origin. Q represents a point which has undergone partial Pb loss,while T0 indicates the original age of a sample which has lost all radiogenic Pb at time T, while all U-Pbsystems should fall on the discordia line joining these two points [from \Faure, 1986 #1347].


0

0.05

0.1

0.15

0.2

0.25

0.3

0.35

200 300 400 500 600 700238 U / 206 Pb

10Ma12142030

Concordia

initial Pb composition

Pb loss

Age of sample

207

Pb /

206

Pb

Figure D-2: Tera-Wasserburg Concordia diagram of measured ratios from a young zircon sample. Thehollow arrow represents a line projecting back to the y intercept that connects the results with an initialPb composition. The solid arrow represents the direction samples should plot if they have undergone Pbloss (or U gain). The short arrow indicates the age of the sample shown by the interesction of a line fittedto the results with Concordia. The lack of scatter of results away from the hollow arrow indicates that Pbloss has not been significant in the analysed zircons. [See \Tera, 1972 #1319].

D3. Secondary Ion Mass Spectrometry

To calculate the age of a sample using U-Pb or Th-Pb decay schemes, it is

necessary to measure the elemental daughter/parent ratios of these elements using a mass

spectrometer. A mass spectrometer is designed to generate ionised species from a sample

and separate these atomic and molecular ions on the basis of their charges and masses,

according to their behaviour in magnetic and/or electric fields. The separated isotopic

species are measured in a collector using an electron multiplier or Faraday cup detector.

An ion microprobe is a type of mass spectrometer that uses secondary ions

produced from the surface of a sample by bombarding it with a beam of energetic primary

ions. The overall process is known as secondary ion mass spectrometry (SIMS). This

form of mass spectrometry allows the in situ isotopic analysis of material, unlike other

methods that require chemical separation of elements before measurement. Secondary

ions are physically eroded from the sample surface by the primary beam of charged

particles (with energies typically on the order of 10 keV), a process known as

“sputtering”. As with other forms of ionisation only a small proportion of the material

ejected from the sample surface is ionised. Ionisation probability can be increased by the

use of a primary beam of highly electronegative (oxygen) or highly electropositive

(caesium) species in the sputtering region, resulting in enhanced emission of positive and

negative ions, respectively.

Appendix D 239

Different elements have different ionisation potentials, so ion yields during

sputtering vary and do not record the true elemental concentrations. Ion yield is also

influenced by matrix effects and instrumental parameters. Anderson and Hinthorne [,

1973 #1321] likened the sputtering environment to a plasma in local thermodynamic

equilibrium (LTE) to develop a model relating secondary ion yields to isotopic abundance.

Although this analogy has been shown to be incorrect, the LTE model has been

successful in modelling the true elemental concentrations.

Mass fractionation between isotopes of the same element is a by-product of the

sputtering process, with lighter isotopes preferentially ionised compared to heavier

isotopes (Slodzian et al., 1980). The amount of isotopic fractionation is proportional to

the fractional mass difference. It is also matrix-dependent and can be affected by

instrumental parameters (Shimizu and Hart, 1982). Instrumental mass fractionation can

be monitored in principle by measuring the isotopic ratios of gravimetric blends of pure

isotopes. However, such standards are only available as pure metals or silicate glasses.

The complex mass spectrum produced during sputtering can include atoms and

molecules of all elements present in both the sample surface and the primary ion beam.

This results in complicated molecular isobaric interferences, especially for intermediate

masses, which require separation by peak stripping or by measurement with sufficiently

high mass resolution to separate different peaks.

D4. Sensitive High Mass Resolution Ion MicroProbe(SHRIMP)

SHRIMP was developed to achieve the high mass resolution necessary for the

removal of isobaric interferences from the isotopic species of interest in U-Th-Pb

geochronology, while retaining high sensitivity (Clement et al., 1977). This was

achieved in two steps. The first was the use of a physically large secondary mass

analyser with a very large magnet-turning-radius relative to “normal” mass spectrometers.

This achieves high mass dispersion which allows use of a wide source slit for maximum

ion transmission. Second, the mass resolution was improved by using an ion optical

system that corrects the second-order focussing aberrations that are normally tolerated in

sector mass spectrometers. A schematic diagram of SHRIMP II is shown in Figure D-3.

The primary ion source is a hollow-cathode duoplasmatron that produces a negative

primary ion beam from oxygen gas through the application of a 450 volt arc discharge

between a Ni cathode and anode plate. The primary beam is extracted through an aperture

in the anode plate by application of a 10 keV accelerating potential to an extraction

electrode. A Wien velocity filter selects primary beam ions of the requisite mass (usually

O2-) and reduces hydride interferences by eliminating species such as OH-. The primary

beam is then demagnified through a series of einzel lenses configured to produce Köhler


illumination. This has the advantage of producing a sharply defined area of sputtering

over which the ion-density of primary ions is uniform.

The primary beam hits the surface of the sample at an incident angle of 45˚

producing elliptical craters that range in size from 10-30 µm in diameter, depending on the

diameter of the Köhler aperture selected. The size of the aperture limits the total ion

current available and is selected by the operator, dependent on whether spatial resolution

or higher secondary ion emission is the more important during analysis. Charge build-up

on the sample surface during sputtering is reduced by the application of a conductive

coating to the sample (generally of 99.999% pure gold). The target is maintained at a 10

keV potential. Positive secondary ions are extracted perpendicular to the sample surface

and accelerated towards an intermediate electrode and extraction aperture at ground

potential. The secondary ion optical array focusses the beam into the entrance slit of the

mass analyser and utilises phase space concepts of beam transport theory to maximise

transmission (Clement et al., 1977). Approximately 10% of the secondary beam is used

as a secondary beam monitor (SBM) to identify instability in the primary beam.

SHRIMP is a double-focusing mass spectrometer that allows angular and energy

refocussing using electrostatic and magnetic sectors to focus into a collector the wide

range of initial energy in the secondary ions produced by the sputtering process. The

configuration is based on a design by Matsuda [, 1974 #112] and utilises a cylindrical 85˚

electrostatic analyser (ESA) (turning radius 1.27m), and a 72.5˚ magnetic sector (turning

radius 1 m), separated by an electrostatic quadrupole lens. SHRIMP operates in single

collector mode with ions measured as pulse counts by a single electron multiplier. The

secondary ion beam intensity is generally insufficient for measurement by Faraday cups

(although these are available for use), which have higher noise levels than electron

multipliers. The measurement of high count rates is affected by dead time and hence a

dead time correction factor must be applied for isotopic ratios associated with high count

rates.

SHRIMP’s mass analyser configuration produces a wide separation of masses at

the collector. A mass resolution of around 5500 is required for U-Th-Pb analysis of

zircon. This resolution separates all masses of interest from their interferences (except Pb

hydrides), so that peak-stripping is unnecessary. A source slit width of 80 µm generates

Pb+ sensitivity in zircon exceeding 20 counts/s/ppm Pb/nA for SHRIMP II, and ranging

between 5-10 counts/s/ppm Pb/nA for SHRIMP I. The magnet is cyclically stepped

through field positions equivalent to the atomic masses to be measured. This process is

controlled by a computer program developed by R. Dabrowski, which on SHRIMP II has

recently been updated to a LabVIEW program by P. Lanc.

Appendix D 241

SHRIMP II

duoplasmatron

Wien filter

Kohler lensKohler apeture

Condensor lens

sampleextractionlens

quadrupoletriplet

source slit

electrostaticanalyser

energy defining slitquadrupolelens

COLLECTOR

collector slitFaraday cup 1melectron multiplier

magneticanalyser

MASS ANALYSER

SOURCE

beam limiting apeture

1270

mm

1000 mm

Figure D-3: Sketch of SHRIMP II design

D5. SHRIMP Data Collection

Table D-2 illustrates a typical mass analysis cycle for zircon U-Th-Pb

geochronology, with the mass analyser stepping from light masses to heavy masses, and

with time allowed for automatic centering of the peaks between measurement of different

species.

Table D-2: Typical meaurement cycle for zircon analysis

Species Nominal Mass (amu) Count Time (s)

Zr2O 196 2

204Pb 204 10

Background 204.1 10206

Pb 206 10207

Pb 207 30208

Pb 208 10238

U 238 5

ThO 248 5

UO 254 2

Mass 196, Zr2O, is used as a reference peak for zircon analyses because its

relatively high abundance aids peak centering during analysis as the magnet cycles from

high mass back to low mass. The 204

Pb peak is too small for autocentering purposes, so

the use of a reference peak with lower mass is very desirable in determining the correct

setting for the magnet. During analysis of titanite, an unidentified peak at mass 200 is


used for this purpose, while in monazite a CePO4 peak at mass 203 is used. The isotopic

species being measured and the count time devoted to each one may be varied by the

operator depending on the type of sample and its age. In general, samples older than

1000 Ma require an accurate measure of 204

Pb to assess common Pb versus radiogenic Pb

contents (see Section D6. D6.4), hence count times on this species need to be increased.

Young samples (< 1000 Ma) rely on a correction for common Pb based on measurement

of 207

Pb or 208

Pb so the count rates on these species are optimised, while 204

Pb may be

neglected altogether.

In a single collector assembly utilising an ion counter such as SHRIMP, the

expected precision of results is governed mainly by Poisson counting uncertainty. This

will be illustrated in Section D6. D8 where the data reduction procedure and estimations

of the uncertainties in the various ion-current ratios are discussed. During on-line data

acquisition, the total count period for each mass is divided into ten time segments, with

outliers identified from Poisson counting statistics being rejected before the mean count

rate is calculated (where the standard deviation of the distribution is the square root of the

mean). This can minimise the effects of short-term beam instabilities affecting individual

measurements for each mass station. Temporal variations in secondary ion emission

(e.g., from sample heterogeneity or primary beam instability) are estimated by cycling the

magnet to measure the mass stations of interest (Table D-2) several times. Usually 5 to 7

cycles are employed (occasionally up to 9 cycles). Any variations in emission are

corrected for by constructing a line of best fit through the data that traces the beam

intensity for each mass over time. Secondary ion ratios are formed from the count rates

for each species at the analysis midpoint using an in-house computer program “PRAWN”

written initially by P. Kinny and refined by T. Ireland (Figure D-4). The scatter of ratios

relative to the fitted line is tested against the scatter expected from ion counting statistics

alone. Ususally the data pass the test but beam instability sometimes causes it to fail, in

which case the error of the mean is set to the empirically-observed scatter. The ratios may

also fail the test if there are changes in Pb/U in the target due to in situ Pb loss or overlap

of the beam onto a mineral growth zone of different age. In addition to this procedure,

the effects of primary beam instability can be corrected by taking the ratio of counts for

each mass to the Secondary Beam Monitor (SBM). This can either be done during

analysis, or during the first stages of data reduction utilising PRAWN.

Appendix D 243Prawn 6.5.3 Output

FAIL

Zr20

PASS

F = 3.1

F = 0.3

PASS

bkgrnd

F = 0.2

PASS

206Pb

F = 2.2

PASS

207Pb

F = 0.8

PASS

208PbF = 2.2

PASS

F = 2.1

PASS

F = 1.6

PASS

UO

F = 2.4

SBM

16000

170005300

5800380

480650

7505

1510

18180

230

0.20

2.003600

4000

238

U

ThO

552346

424.846.53

696.899.55

9.400.79

14.030.32

205.033.77

0.020.03

0.180.26

383143

204Pb

0.00

c / s+

+

+

+

+

+

+

+

+

-0.90

c / s

c / s

c / s

c / s

c / s

c / s

c / s

c / s

c / s

Figure D-4: Output from PRAWN 6.5.3 program showing the counts per second (c/s) for each speciesmeasured, the line fit through the data points for each scan and the c/s for the secondary beam moniotor(SBM).


D6. SHRIMP Data Reduction

D6.1. SHRIMP Standards

A vital aspect of SHRIMP analysis is the use of mineral standards for which the

isotopic compositions have been determined by mass-spectrometric isotope dilution. As

SHRIMP can not measure the absolute concentration of isotopic species, the use of a

standard is necessary to determine elemental Pb/U values of samples of unknown

composition. Ideally, SHRIMP standards must be homogeneous in mineralogy,

crystallinity and chemical and isotopic composition, with concordant Pb/U and Pb/Th

ratios and preferably uniform U and Th contents. Standards should preferably have

sufficiently high uranium contents (hence high radiogenic Pb contents) to produce precise

measurements through low ion-counting errors but not so much uranium that the structure

of the standard is modified by metamictisation. Standards should also contain minimal

common Pb and be readily available. Due to the unquantified influence of matrix effects,

SHRIMP standards must be as close in composition as possible to the unknowns being

analysed. The following minerals and their standards have been used in this study:

D6.1.1. Zircon

A natural gem quality zircon from Sri Lanka, SL13, was used for all analyses.

Based on TIMS analyses listed by Claoue-Long et al. (1995), it has a weighted mean206Pb*/238U of 0.092821 ± 0.000054 (2σ) equivalent to an age of 572.2 ± 0.4 Ma (2σ)

and contains 236 ppm U. Known variations of at least 15% in the U content of SL13

(Ireland, 1995) limits the accuracy of elemental abundances calculated for U, Th and Pb.

Although the Th/U has not been measured by isotope dilution due to the dangers inherent

in using Th spikes, SHRIMP analyses have shown that measured 232ThO+/238UO+ is

directly proportional to target 232Th/238U. The SL13 standard is less uniform in Pb/U

than is desirable, having an external reproducibility of ~ ±2%. Some unknowns have

reproducibilities of ~ ±1%, indicating that the 2% reproducibility of the standard is not an

analytical problem related to sputtering or secondary ion extraction, but reflects

heterogeneity in the composition of the standard. Compston (1996) interprets the excess

scatter in SL13 as due to a bimodal age population within SL13, produced by Pb

redistribution some 15 Ma after its original crystallisation. These effects are averaged out

and not detected by TIMS analyses owing to the much greater sampling volumes

(typically on the order of 30 µg of material or more, whereas SHRIMP typically measures

less than 5 ng). This variation might also be due to occasional µm-scale patches of

unsupported radiogenic Pb (Williams et al., 1984) in SL13.

Appendix D 245

D6.1.2. Titanite

The first reported SHRIMP U-Pb age determinations on titanite were by Kinny et

al. (1994) who characterised a standard using conventional Thermal Ionistaion Mass

Spectrometry (TIMS), supplemented by SHRIMP work at Curtin University. The

standard consists of fragments of titanite collected by N.J. McNaughton from the Khan

pegmatite, Namibia. This standard has been adopted for SHRIMP titanite dating at

A.N.U.. It produces a conventional TIMS age of 518 Ma, has high U contents (~ 696

ppm) and a consistent within-grain composition with moderately uniform Pb/U ratios. A

small degree of Pb loss increases the percentage error on 206Pb/238U ages. This means

that the standard may be used for dating Phanerozoic samples (where its error has only a

relatively small effect on the error of the sample), and also very old samples (> 1500Ma)

because a standard is not required for the derivation of 207Pb/206Pb ages. However, it is

not well-suited for dating intermediate-aged samples. Titanites can be analysed with only

minor adjustments to the runtable commonly employed for zircon analyses on SHRIMP.

An unidentified peak which is always found in titanite at mass 200 is measured (Figure

D-5). This peak has been desrcibed as a CaTiO+ peak (Kinny et al., 1994) but mass

balance calculations suggest the closest corresponding peak would be Ti4+.

0

1

2

3

4

Cou

nts

(100

K)

Mass (AMU)200.0 200.5199.5

Figure D-5: Characteristic peak shape at mass 200 in titanite.

D6.1.3. Monazite

Grains from the Delegate Adamellite, New South Wales, were used as a reference

material. These have yielded an Isotope Dilution Thermal Ion Mass Spectrometry

(IDTIMS) age of 426.7 Ma and a uranium content of 1164 ppm (Williams et al., 1983).

While this age appears to agree with both IDTIMS and SHRIMP ages for zircon from the

same samples, there is a noticeable dispersion in both zircon and monazite ages, which


suggests the presence of more than one age population. Details of the monazite standard

are given in Sircombe [, 1997 #1335]. The reference peak used at low mass is a Ce-

phosphate at mass 203.

D6.2. Hydride Interferences

Pb hydrides (PbH+ species) interfere with the Pb+ mass spectra, mainly acting to

increase measured 207Pb/206Pb ratios as 206PbH+ contributes to the measured 207Pb

peak. If the hydrides are produced by water adsorbed on the sample or in the sample

chamber, the ratio will decrease over time as the vacuum improves, and so the problem

can be identified (Figure D-6).

0.05

0.1

0.15

0.2

0.25

0.3

0.35

0.4

0 5 10 15 20 25

Pb h

ydri

des

Time (hours)

Figure D-6: Plot of Pb hydrides vs time for a zircon sample measured during one analytical session,showing the lack of correlation between Pb hydrides and time. A positive correlation between Pb hydridesand time is expected if hydrides have influenced measurements.

D6.3. Calculation of inter-element ratios

One difficulty in applying SHRIMP to U-Th-Pb geochronology is that, as in all

SIMS analyses, sputtering produces complex mass spectra where each element forms

diverse molecular species reflecting the compositional diversity of the sample and primary

beam. Uranium occurs in the secondary beam as UO2+, UO+ and U+, thorium as

ThO2+, ThO+ and Th+, while lead occurs almost entirely as Pb+. This means that the

ratio of Pb+/U+ in the secondary beam cannot be directly related to the elemental Pb/U

Appendix D 247

ratio in the sample essential for age determinations. (Andersen and Hinthorne, 1972)

showed that measured Pb+/U+ is related to the UO+/U+ measured at the same time. This

led Compston et al. (1984) to employ measured values of UO+/U+ and Pb+/U+.to

calculate Pb/U in the sample.

The relationship between 206Pb+/U+ and UO+/U+ in zircon is approximately linear

(Compston et al., 1984) and can be described as:

( / ) . [( / ) . ]Pb U UO Ustd std+ + + += −0 0764 2 77

Over wide ranges of UO+/U+ some curvature is observed in the relation which has

also been described as quadratic (Williams and Claesson, 1987):

( / ) . ( / ) . ( / ) .Pb U UO U UO Ustd std std+ + + + + += + −0 0048 0 0265 0 08252

Definition of the curve has subsequently been improved using a power law fit

(Claoue-Long et al., 1995):

( / ) . ( / ) .Pb U UO Ustd std+ + + += 0 0069 1 979

The exponent of the power law correlation is given by the slope of the best fit

regression through the logarithms of Pb+/U+ and UO+/U+ for the standard data. For

zircon this slope is measured as 2.00 ± 0.05 (Claoue-Long et al., 1995). The power law

form is regarded as the best expression of the relationship between Pb+/U+ and UO+/U+.

Using the above relationship, the expected Pb+/U+ for the standard can be calculated for

the particular UO+/U+ values for each analysis of the unknowns. For this condition of

equal UO+/U+, it is then assumed that

(Pb / U)unk

(Pb+ / U + )unk

=(Pb / U)std

(Pb+ / U + )std

Because (Pb/U)std is known (approximately) at 0.0928 for SL13, (Pb/U)unk can

now be calculated. For some minerals, such as perovskite, the correlation betweenPb+/U+ and UO+/U+ is not as strong as that between Pb+/UO+ and UO2+/UO+.

A directly proportional relationship has been found between measured232ThO+/238UO+ species and elemental 232Th/238U in the sample (Compston et al.,

1984). These workers defined a constant of proportionality (K=1.11) such that:

232Th238U

= 1.11232ThO+

238UO+


This relation has been modified by Williams (1996) to account for wide-ranging

UO+/U+:

232Th238U

= 0.03446 UO+ / U +( )+0.868[ ]232ThO+

238UO+

Since 208 232 232 1Pb Th et* ( )= −λ

and 206 238 238 1Pb U et* ( )= −λ

it follows that 232

238

208

206

238

232

1

1

Th

U

Pb

Pb

e

e

t

t *

*= ×

−

−

λ

λ

and that the constant of proportionality (K) between 232ThO+ /238UO+ and 232Th/238U can

be defined by the equation

K

Pb

Pb

ThO

UO

e

emeas

t

t

*

*

=

×

−

−

+

+

208

206

232

238

232

238

1

1

λ

λ

Using this equation, values of K have been determined for titanite (this study) as:

232Th238U

= 1.0213232ThO+

238UO+

and for monazite (Sircombe, 1997):

232Th238U

= 0.8437232ThO+

238UO+

The calibration of Pb+/ThO+ and UO+/U+ can also be determined from the slope of

the best fit regression through the logarithms of these ratios from the standard. Slopes

determined by this method include: 1.00 ± 0.1 for zircon, while the slope for titanite has

been calculated as 0.88 ± 0.09 (Table D-3). A slope of 2.3 ± 0.5 has been calculated

from monazite for a regression line of ln(Pb+/Th+) vs ln(UO+/U+), rather than

ln(Pb+/ThO+) vs ln(UO+/U+) (Sircombe, 1997). The difference in the slopes for

ln(Pb+/ThO+) compared to ln(Pb+/Th+) can be explained according to the following

equations:

Appendix D 249

ln ln lnPb

ThA m

UO

U

= +

+

+

ln ln ln lnPb

ThO

Pb

ThUO

U

Pb

Th

UO

U

=

=

−

+

+

+

+

ln ln ln lnPb

ThOA m

UO

U

UO

U

= +

−

+

+

+

+

ln ln lnPb

ThOA m

UO

U

= + −( ) ×

+

+1

where m is the slope of the calibration line for Pb/Th and ln(A) is the intercept. For

Pb/ThO ratios the slope is m-1, so if the calibration slope for ln(Pb/Th) is 2.0 then the

slope for ln(Pb/ThO) will be 1.0.

Table D-3: Calibration slopes from 11 sessions on the titanite standard Khan Pegmatite(Z2265 excluded)

Sample No.

analyses

U/Pb slope Th/Pb slope lnU+/CaTi2O4

Z1800a 22 2.58 ± 0.22 1.75 ± 0.28 -2.56

Z1800b 14 1.58 ± 0.09 -0.10 ± 0.32 -2.02

Z1800c 33 2.53 ± 0.40 -0.20 ± 0.84 -1.46

Z2155 16 1.73 ± 0.14 0.88 ± 0.14 -2.97

Z2217 15 1.63 ± 0.23 0.69 ± 0.31 -1.48

Z2241 13 1.60 ± 0.20 0.79 ± 0.27 -2.24

Z2265 16 0.74 ± 0.98 -8.55 ± 17.1 -1.41

Z2313 20 1.59 ± 0.13 0.58 ± 0.20 -2.54

Z2362 11 1.81 ± 0.12 1.10 ± 0.58 -3.14

Z2600 10 1.95 ± 0.26 1.22 ± 0.31 -2.31

Z2615 7 1.94 ± 0.32 1.13 ± 0.30 -2.33

Results 151 1.72 ± 0.06* 0.88 ± 0.09# -2.31

* weighted mean with MSWD 3.0

# weighted mean with MSWD 1.7

Elemental Abundances

Abundances of U, Th and Pb are calculated empirically in zircon by relating

uranium concentration in the sample to Zr content at the same values for UO+/U+

(Compston et al., 1984):


(Zr2O / U)unk

(Zr2O+ / U + )unk

=(Zr2O / U)std

(Zr2O+ / U + )std

Zr2O+/U+ in the standard is adjusted to be for the same UO+/U+ as the unknown

using a power law curve (Claoue-Long et al., 1995):

Zr2O+

U + = aUO+

U +

0.66

where “a” is determined from the mean of meaurements from the standard over an

individual analytical session, and 0.66 is the slope of the calibration line of ln(Zr+/U+)

versus ln(UO+/U+).

Uranium concentrations are derived from the expression:

Uconc UconcZr O U

Zr O Uunk stdstd

unk

=

+ +

+ +

( / )

( / )2

2

where the uranium concentration of the zircon standard SL13 is taken as 236 ppm,

as measured by isotope dilution.

Once U concentrations are established, Th and Pb abundances can be calculated

from Th/U and Pb/U ratios.

This procedure works well for samples with little variation in matrix characteristics

such as titanite, where a similar relation between the measured Ti+ reference peak and

UO+/U+ has been identified.

ln .

Ti

Ua

UO

U4

2 31+

+

+

+

−

=

Uconc UconcTi U

Ti Uunk stdstd

unk

( / )

( / )=

+ +

+ +4

4

where -2.31 is the slope of the calibration line of ln(U+/Ti4+) versus ln(UO+/U+)

and the Khan pegmatite titanite standard contains approximately 697 ppm uranium from

isotope dilution measurements (Fanning, unpubl.).

Appendix D 251

For other minerals such as monazite, which has variable composition due to

substitution of different elements (Sircombe, 1997), such a relationship cannot be

established. Instead, the U ppm can be approximated from the number of UO+ counts per

second, assuming the material measured as a standard has the same content of uranium as

that measured by isotope dilution.

Uppm UppmUO cps

mean UO cpsunk stdunk

std

=

+

+

where the Delegate adamellite monazite contains 1164 ppm uranium (Williams et

al., 1983).

D6.4. Common Pb corrections

Lead consists of four different isotopes 204Pb, 206Pb, 207Pb and 208Pb. Only204Pb is not produced by radioactive decay. In order to determine an absolute age using

U-Th-Pb geochronology, it is necessary to be able to separate the component of common

Pb incorporated into the crystal during or following recrystallisation from the amount of

radiogenic Pb. Total common Pb in an analysis is the sum of: 1) initial common Pb in the

sample, both common Pb incorporated in mineral lattice and in submicron-scale mineral

inclusions; 2) Pb incorporated by post-crystallisation exchange with surrounding material

that has evolved beyond initial compositions; 3) surface contaminants. The effect of

surface contaminants is reduced during SHRIMP analyses by rastering the target area

with the primary beam for 3 minutes before analysis. Uncertainty in common Pb can be

reduced by thorough cleaning of samples to remove surface Pb contaminants before

analysis and by careful selection of the beam target to avoid cracks and inclusions. There

are three different ways of estimating the proportion of common Pb in a sample - using

the measured amounts of 204Pb, 207Pb and 208Pb respectively.

D6.4.1. 204Pb Correction

If the initial Pb composition is known, then the amount of common Pb is directly

proportional to the amount of 204Pb. The atomic fraction of 206Pb which is common can

be defined as:

f = 206Pbcommon/ 206Pbmeas

As all measured 204Pb is common, the fraction of common 206Pb can be defined as:


f 4 =204Pb/ 206Pbmeas

204Pb/ 206Pbcommon

where 204Pb/206Pbcommon is assumed or known (see Section D6.5).

The 204Pb correction is often used, especially for old samples to calculate207Pb/206Pb ages. The main disadvantage is that it is the least abundant Pb isotope and

thus has large uncertainties due to low count rates and also requires a significant amount

of available analysis time. It is also subject to isobaric interferences such as 186W18O,204Hg and in monazite an unidentified molecular species (Sircombe, 1997). This makes

the correction unsuitable for young zircons, which rely on small uncertainties in the

amount of common Pb.


This method is used exclusively for young ages (< 800 Ma) as the long half-life and

small relative abundance of 235U results in little radiogenic 207Pb accumulation in such

samples, so that the radiogenic 207Pb/206Pb ratio for the sample can be closely estimated.

This method is only usable for determining 206Pb/238U ratios, as use of this correction

assumes a value for 207Pb/206Pb ratios.

207 206235

238

235

238

1

1Pb Pb

U

U

e

e

t

t* */ =

−

−

λ

λ

fPb Pb Pb Pb

Pb Pb Pb Pbmeas meas

com com7

207 206 207 206

207 206 207 206=( ) − ( )( ) −( )

/ /

/ /

* *

* *

By assuming Pb/U concordance, this method is equivalent to extrapolating each

analysis along the common Pb mixing line on a Tera-Wasserburg Concordia diagram until

the Concordia curve itself is intersected. The corrected 206Pb*/238U value is then

obtained from the Concordia x-axis.


This correction relies on radiogenic 208Pb*/206Pb* being estimated from ThO+/U+

for an assumed formation age. As the correction is largely insensitive to the choice of

formation age it can be applied to samples of any age.

208 206232

238

232

238

1

1Pb Pb

Th

U

e

e

t

t* */ =

−

−

λ

λ

Appendix D 253

fPb Pb Pb Pb

Pb Pb Pb Pbmeas meas

com com8

208 206 208 206

208 206 208 206=( ) − ( )( ) −( )

/ /

/ /

* *

* *

This correction is only suitable for samples which have low Th/U ratios, which will

accumulate only small amounts of radiogenic 208Pb from decay of 232Th. The 204Pb

correction method is favoured over the 208Pb for determining old (>800 Ma) as old

samples may have accumulated a large proportion of radiogenic 208Pb. As the 208Pb

correction relies on the assumption of closed system evolution of Th/U ratios, differential

movement of U and Th atoms in minerals and the preferential loss of 208Pb can invalidate

its use in estimation of the common Pb content of zircons. To assess whether it is valid to

apply the 208Pb correction for a set of zircon analyses, a plot of radiogenic 208Pb/206Pb

(calculated independently using the 204Pb or 207Pb methods) versus 232Th/238U

(Compston et al., 1984) should yield a line to the origin. The slope of this line is

relatively constant for systems of all ages (Figure D-7) as the half-lives of 232Th and238U are similar in magnitude.

Th/U

0.3

0.2

0.1

0.00.0 0.2 0.4 0.6 0.8 1.0 1.2

208

Pb*

/ 206

Pb*

Figure D-7: To validate the use of the 208-corrected method in calculating radiogenic Pb, data shouldshow a straight line relationship between 207-corrected 208Pb*/206Pb* values and Th/U, as shown inthis illustration.

D6.5. Common Pb composition

In minerals such as titanite, which may incorporate significant amounts of common

Pb in their crystal lattice, the selection of common Pb composition can be critical. The

common Pb isotopic composition can be assessed by several methods: it can be directly


measured from minerals which incorporate Pb in their structures but exclude U, such as

feldspars and galena or it can be independently calculated using measured isotopic ratios

from the mineral being dated. The common 204Pb composition can be calculated from

from the intercept of 204Pb/206Pb versus 238U/206Pb, while the common 207Pb

composition can be calculated from the intercept of 207Pb/206Pb versus 238U/206Pb

(Figure D-8), according to the equations:

204

206

204

206

204

206

238

206238 1

Pb

Pb

Pb

Pb

Pb

Pb

U

Pbe

m c c m

t

=

−

×

× −( ) λ

207

206

207

206

207

206

206

238

238

206

207

206

Pb

Pb

Pb

Pb

Pb

Pb

Pb

U

U

Pb

Pb

Pbm c m m c

=

−

×

×

+

*

*

*

where m stands for measured value, c stands for common value and * represents

the radiogenic value.

The common 208Pb component can be evaluated once the common 204Pb or 207Pb

compositions are known. Using the appropriate f4 or f7 value from the above equations,

the common 208Pb can be calculated from radiogenic 208Pb/206Pb and Th/U according to

the following equation:

208

206

208

206

208

2061Pb

Pbf

Pb

Pb

Pb

Pbf

m c

− −( ) ×

=

×

*

10 Ma15

207

Pb /

206

Pb

238 U / 206 Pb

0.0

0.2

0.4

0.6

0.8

1.0

0 100 200 300 400 500 600 700

common Pb composition

age of sample

Concordia

Figure D-8: Line fit through measured isotope ratios to determine the common Pb composition fortitanites from a particular sample.

Appendix D 255

If the mineral being analysed has low initial common Pb values, ie. Pb is

preferentially excluded from the crystal structure (zircon, monazite), then the isotopic

composition of common Pb must be calculated according to the age of the sample using a

model for terrestrial Pb isotopic evolution over time.

Patterson (1956) measured the isotopic composition of Pb in meteorites and

calculated their age as 4.55 ± 0.07 Ga (207Pb/206Pb method), defining the isotopic

evolution of Pb from this starting point. As the composition of some Pb isotope samples

from Earth also lies close to the isochron for meteorites Patterson suggested that this was

also the age of the Earth and that the isotopic evolution of terrestrial Pb followed the same

single stage evolution as that of meteorites. This could then be defined according to the

single-stage equations of the Holmes-Houtermans model where:

206

204

206

204238 238Pb

Pb

Pb

Pbe e

c i

T t

=

+ × −( )× ×µ λ λ

207

204

207

204

235

238235 235Pb

Pb

Pb

Pb

U

Ue e

c i

T t

=

+ × × −( )× ×µ λ λ

208

204

208

204

232

238232 232Pb

Pb

Pb

Pb

Th

Ue e

c i

T t

=

+ × × −( )× ×µ λ λ

where c and i stand for common and initial ratios, respectively, T is the age of the

Earth, t is the age of the sample and µ is the 238U/204Pb.

Further analysis of samples of terrestrial Pb revealed a paradox however, where the

Pb isotopic composition of present day rocks is incompatible with simple closed-system

evolution over 4.55 Ga (Doe and Stacey, 1974; Oversby, 1974; Vollmer, 1977)

suggesting 206Pb/204Pb enrichment of the bulk Earth (Allegre et al., 1982). This

discrepancy is thought to be caused by extraction of Pb preferentially over U into the

Earth's core during accretion. Some workers suggest that the Earth experienced no

significant isotopic growth in the time betwen formation of the solar system (~4.56 Ga)

and the end of accretion which is estimated to have occurred over 80 ± 40 Ma (Galer and

Goldstein, 1996). If this stage was one of low µ values then a single-stage Pb growth

model can be used to describe the evolution of terrestrial Pb from the time at which

accretion ceased. Manhes et al. [, 1979 #1147] suggest this occurred at 4.49 ± 0.17 Ga,

in this case this value can be substituted for T. Other workers favour more complicated

Pb evolution models (Sinha and Tilton, 1973; Stacey and Kramers, 1975) and suggest


that continental crust formation, which occurred for at least 400 Ma after Earth accretion,

has an effect on the growth curves .

Common Pb composition for most of the analyses reported in this thesis are based

on the measured value of Pb isotopic composition for troilite from the Canyon Diablo

meteorite (Tatsumoto et al., 1973), with the isotopic evolution of crustal lead modelled as

a single stage process from 4.5 Ga (Manhes et al., 1979). Accordingly the following

expressions can be derived:

204

206 238 4500 238

1

9 307

Pb

Pb e ect

=

+ −( )( )× ×. µ λ λ

207

206235 4500 235

204

20610 294137 88

Pb

Pbe e

Pb

Pbc

t

c

= + × −( ) ×

× ×..

µ λ λ

208

206232 4500 232

204

20629 476 3 8 137 88Pb

Pbe e

Pb

Pbc

t

c

= + × × × −( )( )( ) ×

× ×. . .µ λ λ

where 9.307, 10.294 and 29.476 are the Pb isotopic compositions for the Earth at

formation, as derived from the Canyon Diablo meteorite, 3.8 is the modern 232Th/238U

ratio (Cumming and Richards, 1975) and µ is the modern 238U/204Pb ratio of 8.8 (Allegre

and Lewin, 1989). Total common Pb in multiple zircon populations or for zircons with

complex histories might consist of several different common Pb components. Unusual

common Pb compositions can be allowed for as long as a simple mixing relationship

exists between the common Pb and the radiogenic Pb.

D6.6. Calculation of Radiogenic Isotope Ratios

The proportion of common Pb which a mineral incorporates (f) is used to calculate

the amount of radiogenic Pb from measured isotopic ratios. The calculation of the

radiogenic 206Pb/238U ratio (206Pb*/238U) can be made according to the following steps:

206

238

206

2381

Pb

Uf

Pb

Umeas

*

( )

= −

Radiogenic 207Pb/206Pb (207Pb*/206Pb*) can be calculated from:

207

206

207

206

207

206

1

Pb

Pb

Pb

Pbf

Pb

Pb

fmeas c

*

*

( )

=

−

−

Appendix D 257

208Pb*/206Pb* can be calculated in the same way using f and the measured and

common 208Pb/206Pb ratios. The radiogenic 207Pb/235U ratio can be calculated by

combining 206Pb*/238U and 207Pb*/206Pb* ratios:

207

235

207

206

206

238137 88

Pb

U

Pb

Pb

Pb

U

* *

*

*

.

=

×

×

To calculate 206Pb*/238U for the unknowns, a “calibration” line per analytical

session is fitted to analyses of the standard in terms of the logarithms of 206Pb+/238U+ and

UO+/U+, as described in Section D6.3. Individual analyses for the unknowns are

calculated by comparing the observed values of 206Pb*+/U+ for the unknowns with the

values of 206Pb*+/U+ for the standard expected from that calibration line at the various

UO+/U+ of the unknowns.

206

238

206

238

206

238

Pb

U

Pb

U

Pb

U

**

*

=

×

+

+

+

+

observed

expected

known ratio of standardof unknown

of standard

The individual 206Pb*/238U ratios of the standards per session may be calculated in

the same way. Ages may then be calculated according to the equations:

AgePb

U

Pb

U206

238

206

238

238

1

=

+

ln*

λ

AgePb

Th

Pb

Th208

232

208

232

232

1

=

+

ln*

λ

AgePb

U

Pb

U207

235

207

235

235

1

=

+

ln*

λ


D7. Isotopic Disequilibrium

U and Th decay to Pb via a series of shorter-lived intermediate daughter products

forming a decay chain (Figure D-9). All U-Th-Pb age calculations assume that these

decay chains were in secular equilibrium when radioactive decay started. Secular

equilibrium occurs when there are equal numbers of decays per unit time for all isotopes

in the decay chains. This means that any radioactive daughter isotopes must be initially

present in the mineral in amounts inversely proportional to their decay constants.

Incorporation or exclusion of intermediate daughter products during crystallisation of a

mineral (controlled by partition coefficients and the availability of elements in the

environment of crystallisation) may create disequilibrium in the decay chains. The effects

of isotope disequilibrium are negligble for samples older than ~ 100 Ma where a

sufficiently large amount of radiogenic Pb has accumulated to effectively mask any effect

of initial isotope disequilibrium (Ludwig, 1977; Mattinson, 1973). However, variations

from secular equilibrium can affect the U-Pb systematics of young ( < 100 Ma) samples.

This property is used to good effect in dating very young samples ~ 100 ka (see Faure

(1986) for discussion). Only those intermediate daughter products with long half-lives

will have a significant effect on age calculations.

208Pb/232Th ages can be a good independent geochronometer to compare with the

206Pb/238U ages, as 208Pb/232Th ages are not thought to be affected by isotopic

disequilibrium which may influence young (< 50 Ma) 206Pb/238U ages. For the 232Th-

208Pb decay chain, the longest-lived intermediate daughter has a half-life of only 5.76

days (228Ra) and so isotopic disequilibrium will have a negligible effect on 208Pb/232Th

ages in the range of this study (> 1Ma) (Figure D-9). The 238U-206Pb and 235U-

207Pb decay chains however, have several relatively long-lived intermediate daughter

products, 234U (l=244 ka) and 230Th (l=77 ka) in the 238U-206Pb chain (Figure D-

10)and 231Pa (l=32.5 ka) in the 235U-207Pb chain (Figure D-11).

Appendix D 259

212Po

212Bi

212Pb

216Po

220Rn

224Ra

228Th

228Ac

208Pb

208Tl

228Ra

232Th

126 128 130 132 134 136 138 140 142

90

88

86

84

82

80

Ato

mic

Num

ber

Neutron Number

Figure D-9: Decay scheme for 232Th (from Faure, 1986).

206Pb

206Tl

206Hg

210Po

210Bi

210Pb

210Tl

214Po

214Bi

214Pb

218Rn

218At

218Po

222Rn

226Ra

230Th

234U

234Pa

234Th

238U

92

90

88

86

84

82

80

124 126 128 130 132 134 136 138 140 142 144 146

Neutron Number

Ato

mic

Num

ber

Figure D-10: Decay scheme for 238U (from Faure, 1986).


211Po

211Bi

211Pb

215At

215Po

215Bi

219Rn

219At

223Ra

223Rn

227Th

227Ac

231Pa

231Th

235U

207Pb

207Tl

126 128 130 132 134 136 138 140 142

90

88

86

84

82

80

92

144124

Neutron Number

Ato

mic

Num

ber

Figure D-11: Decay scheme for 235U (from Faure, 1986).

Because normal chemical processes will not fractionate 234U relative to 238U, 234U

can generally be assumed to be in secular equilibrium with 238U and hence ignored from

calculations for isotope disequilibrium (Mattinson, 1973) (although see Ludwig (1977)

for a discussion of the activity ratio of 234U/238U in uranium ore bodies). As far as 230Th

disequlibrium is concerned, the amount of 230Th in a sample depends on the partition

coefficients of the mineral with respect to Th. In general Th is strongly depleted relative

to uranium in both zircon and titanite, while it is preferentially incorporated into the

monazite structure. This leads to a potential deficit of 230Th in zircon and titanite which

could result in a deficit in the amount of radiogenic 206Pb produced by decay along the238U-206Pb chain and a potential excess of 230Th in monazite resulting in excess

radiogenic 206Pb being produced. This would translate into anomalously old 206Pb/238U

ages for monazite, and anomalously young ages for zircon and titanite.

The following correction can be applied to account for excess or deficit 206Pb

produced by 230Th disequilibrium in igneous rock. Assuming that the measured Th/U of

the rock reflects the amount of Th and U available to the crystallising mineral phase, the

amount of excess/deficit 230Th can be calculated from the apparent degree of fractionation

between mineral and magma:

Appendix D 261

FTh U

Th Ueral

magma

=( )( )

/

/min

This can be used to calculate the “real” 206Pb/238U age from the following equation:

206

238

206

238238

2301

Pb

U

Pb

UF

corr meas

=

− −( )

λλ

which should result in the corrected ratios plotting along Concordia (Scharer, 1984).

(Parrish, 1990) described the amount of excess/deficit 206Pb as related to F according to:

t Fage± =

+ −( )

11 1

238

238

230λλλ

ln

The correlation between excess/deficit amounts of 206Pb can be seen in Figure D-12.

0

-10%

-20%

+10%

+20%

+30%

+40%

+50%

206Pb

10 20 30 40 50 60

monaziteallanite

zircon

f=40

f = 20f = 10

f =

Th

U

xenotime

ThU

f = 0.1

f = 1

rock

mineral

(magma)

defi

cit

exce

ss

21.9

Ma

24.0

Ma

(Ma)age

Figure D-12: Correlation between the relative excess/deficit amounts of 206Pb* in responseto excess or deficit 230Th, shown in relation to Th/U fractionation factors, [from Figure 4 in \Scharer,1984 #1331].

These corrections assume that measured whole-rock Th/U represents the effective

Th/U available at the time the mineral crystallised. This assumption may not be valid for

metamorphic minerals, hydrothermally-precipitated minerals or minerals crystallised after

early growth of another mineral phase enriched in Th or U (eg. thorite, allanite, monazite,


zircon), and does not allow for the presence of restitic Th or U-rich minerals which will

influence the whole rock values. The only certain way to allow for 230Th disequilibrium

is to disregard 206Pb/238U ages in favour of 208Pb/232Th ages which are unaffected by

disequilibrium from intermediate daughter products in the decay chain. In general the

effects of a deficit of 206Pb on zircon and titanite ages for samples in the age range of this

study (>12 Ma) are considered to be smaller than the error associated with individual

analyses.231Pa disequilibrium can occur by depletion or enrichment of 231Pa relative to 235U

in a crystallising mineral which would result in a depletion or enrichment in radiogenic207Pb. In zircons, Th, U and Pa substitute for Zr4+, where the ionic radii of these

species in 8-fold co-ordination can be ranked Zr4+ < Pa5+ < Pa4+ < U4+ < Th4+ = 0.84 <

0.91 < 1.00 < 1.01 < 1.05Å (Barth et al., 1994; Berger, 1991). As the ionic radius of

Pa4+ is close to that of U4+ it may be depleted relative to uranium in some minerals,

though not as depleted as Th4+ (Mattinson, 1973). If Pa occurred as Pa5+ then its ionic

radius would be less than that of U4+, resulting in a possible enrichment of Pa relative to

U (Barth et al., 1994). As the geochemical behaviour of Pa is still not well understood, it

is difficult to assess the amount of excess or deficit which may exist (Barth et al., 1994;

Parrish, 1990; Scharer, 1984). Depletion or enrichment of 231Pa will cause depletion or

enrichment of 207Pb which will mainly affect 207Pb/206Pb or 207Pb/235U ages which are

not used in this study (except in samples older than 1000 Ma where the effects of isotopic

disequilibrium will be negligble). Because the measured 207Pb is used to calculate the

proportion of common Pb in most samples younger than 1000 Ma, it could potentially

influence the ages by affecting the amount of correction applied. In samples with low Th

contents 208Pb was also used to calculate the fraction of common Pb in the mineral. All206Pb/238U ages were calculated using the 207- and 208-corrected methods described in

Section D 6.4, and found to be in good agreement.

D8. SHRIMP Error Analysis

The uncertainty quoted on the age of individual SHRIMP spot measurements is a

combination of several factors, including:

1) uncertainties in secondary ion yield as predicted from counting statistics and as

monitored by the quality of fit of a regression line through measurements of isotopic

species over time,

2) the uncertainty in the common Pb correction, controlled mainly by (1)

3) uncertainty in the reference standard calibration line, usually the principal control on the

accuracy of 206Pb/238U measurements.

Appendix D 263

The calculation of uncertainties in secondary ion yield are described in Sections D

D5 and D D6 and these are propagated through subsequent age calculations according to

the following statistical procedure (Bevington and Robinson, 1992):

To determine the error associated with dependent variables (x) that are functions of

one or more different measured variables (a. b, ...), we can expand these measured values

in a Taylor series to form the error propagation equation.

If x = f (a,b,....) where x is a function of the measured quantities independent

variables) a, b, ... and these variables have associated errors (variances) then

σ σ∂∂

σ∂∂

σ σ∂∂

∂∂x a b a b

x

a

x

b

x

a

x

b2 2

22

22 2

2 2

2≅

+

+ +

+ ..... .....

where σ x2 , σa

2 , σb2 are the variances in x, a and b respectively .

This equation neglects the fact that partial derivatives are not constant over the entire

range of variances for the independent variables. The first two terms of the error

propagation equation dominate the uncertainties calculated for x, representing the

uncertainties in a and b weighted by the squares of the partial derivatives ∂∂

∂∂

x

a

x

b ,

.

The third term, which describes the covariance, will usually approximate zero for a

large random sample set, if a and b are uncorrelated. The error propagation equation can

hence be simplified to:

σ σ∂∂

σ∂∂x a b

x

a

x

b2 2

22

2

=

+

+ .....

The standard deviation (σ) of x can then be calculated as:

σ σ∂∂

σ∂∂x a b

x

a

x

b=

+

+ .....2

22

2

This combines all the contributions to the error in quadrature.

This procedure provides an estimate of the precision with which interelement ratios

and trace element concentration have been determined. The accuracy of these values

depends on the uncertainty inherent in the construction of the standard calibration line (see

Section D 6.3). This uncertainty is added in quadrature, as a coefficient of variation, to

the final precision estimates for mean sample Pb/U and Pb/Th and is the main control on

measurement accuracy (Claoue-Long et al., 1995). Analytical uncertainties are reduced

when it is possible to pool several SHRIMP analyses. This is done objectively using the

principles of maximum likelihood theory described in Section D 9.

The error on individual age estimates is quoted at the 1 σ level and the standard

deviation on the mean of age populations is quoted for age groupings.


D8.1. Error on Titanite Ages

Special consideration to the calculation of errors on the age of titanite samples must

be made to account for the large common Pb corrections required. The projection of a

line through the sample ages onto Concordia has a parabolic error envelope (Figure D-8)

which increases the error on the intersection for samples with high comon Pb corrections.

To account for this effect individual titanite analyses and their associated errors were

processed using Ludwig’s ISOPLOT program [Ludwig, 1993 #1426] which calculates

the error on the Concordia intersection. All titanite ages quoted have their error calculated

in this way.

D9. Mixture Modelling

Mixture modelling provides a way of finding the best fit (maximum likelihood) set

of ages and proportions for any number of assumed components (Galbraith and Green,

1990). Using a Gaussian spectral deconvolution program (Sambridge and Compston,

1994) analyses are checked for the existence of multiple age components. This program

objectively searches for “best fit” ages and proportions which will maximise the

likelihood function. This is done by calculating a misfit parameter that rapidly decreases

as the optimum number of distinct components is reached and then slowly decreases once

the optimum is exceeded.

This procedure is applied first to analyses of the standard data where multiple age

components have been recognised and may require a correction factor to analyses of the

unknowns (Compston, 1996). It is then applied to analyses of the unknowns to resolve

the existence of multiple populations.

D10. Statistical Tests on Age data

D10.1. Test of Adequacy

In provenance studies the probability of measuring all significant age components in

a sample can be assessed using a test of adequacy described by (Dodson et al., 1988)

using the equation

P f n= −( )1

where P is the probability of missing a component, f is the proportion of that

component amongst total components and n is the number of randomly selected grains

analysed.

Appendix D 265

D10.2. Significant Differences in Ages

To assess whether two ages are essentially the same or if there is a significant

difference between them requires a significance test.

D a b= + ×σ σ2 2 2 3.

Where a and b are the ages and D is the difference required between a and b for

them to be considered significantly different at the 99% confidence level, and σa and σb

are the errors on the ages (Rowntree, 1981).

D11. KaleidaGraphTM Programs

KaleidaGraphTM formula script were used to assess radiogenic isotope values for206Pb, 207Pb and 208Pb and compare them to standard measurements to determine206Pb/238U, 207Pb/206Pb and 208Pb/232Th ages for unknowns and also Th and U

concentrations according to the procedures outlined in the sections above. 207Pb/206Pb

calculations utililsed a KaleidaGraphTM macro written by Sircombe [, 1997 #1335].

Standard and unknown data from the PRAWN program (version 6.5.5) are run through

separate programs, with unknowns utilising information derived from the standard

results. Copies of the KaleidaGraphTM programs used to calculate the ages presented in

this thesis are available on request.

Appendix E 271

E. : U-TH-PB ANALYTICAL RESULTS

Following are keys to the codes used in the data tables describing the SHRIMP U-Th-

Pb results for each sample.

Correction (Corr) CodeThis column describes the type of common Pb correction applied to each analysis.

1 - 206Pb/238U age corrected using 207Pb

2 - 207Pb/206Pb age corrected using 208Pb

3 - 207Pb/206Pb age corrected using 204Pb

Note that count times were tailored for the measurement of 206Pb/238U as opposed to207Pb/206Pb ratios and count times for 204Pb were often minimised and this is reflected in the

occassionally large errors on 207Pb/206Pb ages. When 204Pb was measured for insufficient

time, the 208Pb correction was applied to calculate 207Pb/206Pb ages.

GRAIN AREA CODE (AREA)This describes the position of the ion probe pit in

relation to the general morphology of the zircon grain

analysed according to the sketch in Figure 1 where:

centre - centre

edge - edge

rim - rim

core - core

inner rim - inrim

core edge - co edge

termination - term.

Note: the ion beam was often positioned to

overlap the plastic resin enclosing the zircon, to ensure

that no overlap of the pit occurred across growth bands.

Figure 1: Sketch of zircon grain displaying the origin of

the descriptive terms for zircon morphology used in this appendix.

20 µm

centre

edge

rim

core

inrim

coedge

term

edge

272 U-Pb Analytical Results

ZIRCON DESCRIPTION CODE (TYPE)

This code gives a brief description of the type of internal zircon structure analysed.

UXC - Unzoned xenocrystic core

OXC - Oscillatory zoned xenocrystic core

IXC - Irregularly zoned xenocrystic core

OZ - Oscillatory zoned grain (sometimes with sector zoning)

IZ - Irregularly zoned grain

OO - Oscillatory zoned overgrowth

UOL - Unzoned overgrowth low luminescence

UOS - Unzoned overgrowth strong luminescence

LOL - Latest overgrowth low luminescence

LOS - Latest overgrowth strong luminescence

SF - Seam infilling fracture

NS - Non-luminescent seam surrounding core

LM - Low luminescent mantle surrounding core

RZ - Recrystallised zone

MZ - Mottled, inclusion rich zone altering pre-existing structures

MZO - Mottled, inclusion rich zone forming overgrowth

MIX - Overlap of different zones

SZ - Sector-zoned grain

AD - Abraded detrital grain

Note: f% is the percentage common Pb for each analysis.

Appendix E 273

E.1 Zircon U-Pb Analytical Results

E.1.1 IO9403 Ios Orthogneiss (Z1978, 97759)

Spot U

(ppm)

Th/U f % 206Pb /238U

±1σσσσ 207Pb /206Pb

±1σσσσ Age

(Ma)

±1σσσσ Corr Area Type

1.1 308 0.43 0.19 0.11052 0.00116 0.05393 0.00042 299.1 3.3 1 in rim oz

2.1 336 0.40 0.11 0.11603 0.00072 0.05353 0.00050 310.0 2.2 1 edge oz

3.1 286 0.45 0.06 0.11048 0.00087 0.05308 0.00087 308.9 2.5 1 in rim oz

4.1 273 0.19 0.75 0.17015 0.00135 0.06324 0.00087 486.4 3.8 1 in rim oz

5.1 258 0.06 0.18 0.10515 0.00103 0.05403 0.00061 305.2 3.0 1 centre oz

5.2 247 0.05 0.01 0.10687 0.00092 0.05272 0.00052 311.7 2.7 1 in rim oz

6.1 103 0.41 0.32 0.21208 0.00128 0.06254 0.00069 597.6 3.6 1 core oxc

6.2 236 0.02 0.15 0.10089 0.00064 0.05382 0.00053 308.2 1.9 1 rim oz

7.1 140 0.51 0.32 0.10711 0.00092 0.05525 0.00076 307.8 2.7 1 rim oz

8.1 319 0.27 0.21 0.15274 0.00135 0.05681 0.00081 413.0 3.8 1 core ixc

9.1 263 0.63 0.21 0.10142 0.00073 0.05425 0.00058 304.6 2.2 1 centre sz

10.1 717 1.34 2.70 0.55361 0.00366 0.11183 0.00080 1090.3 40.0 2 centre oz

11.1 338 0.65 0.57 0.16152 0.00092 0.06139 0.00047 474.5 2.7 1 centre oz

12.1 158 0.45 0.36 0.21176 0.00268 0.06423 0.00105 647.9 7.8 1 centre oz

13.1 263 0.45 0.11 0.10880 0.00052 0.05357 0.00045 314.0 1.6 1 in rim oz

14.1 127 0.38 0.50 0.09949 0.00100 0.05684 0.00064 308.9 3.0 1 core oxc

15.1 258 0.99 0.26 0.21607 0.00286 0.06045 0.00067 540.5 7.4 1 core oxc

16.1 300 0.59 0.19 0.12950 0.00203 0.05418 0.00072 310.0 5.2 1 core oxc

17.1 140 0.44 0.26 0.10709 0.00082 0.05471 0.00060 307.1 2.4 1 in rim oz

18.1 195 1.14 0.74 0.10866 0.00082 0.05879 0.00314 304.6 2.4 1 centre oz

19.1 210 0.36 0.14 0.10475 0.00119 0.05361 0.00063 304.2 3.4 1 centre oz

20.1 486 0.32 1.08 0.10725 0.00101 0.06170 0.00034 305.0 2.9 1 centre oz

21.1 278 0.46 0.11 0.10853 0.00040 0.05340 0.00038 306.0 1.4 1 term oz

22.1 251 0.74 0.91 0.10141 0.00080 0.05997 0.00098 292.0 2.3 1 in rim oz

23.1 260 0.38 0.08 0.10977 0.00062 0.05323 0.00050 309.7 1.9 1 in rim iz

24.1 251 0.15 0.39 0.17407 0.00200 0.05949 0.00062 460.7 5.4 1 centre oz


E.1.2 IO9404 Ios Orthogneiss (Z1978, 97760)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


1.1 193 0.63 0.11 0.10601 0.00061 0.05357 0.00058 312.3 1.8 1 centre oz

2.1 239 0.64 0.09 0.23195 0.00197 0.06244 0.00037 664.5 5.5 1 core oxc

3.1 447 0.26 0.09 0.22455 0.00157 0.06154 0.00035 633.0 4.4 1 core oxc

4.1 468 0.05 0.13 0.11637 0.00039 0.05417 0.00041 331.5 1.3 1 term oz

5.1 128 0.26 0.08 0.09705 0.00066 0.05314 0.00062 305.9 2.0 1 rim oz

6.1 338 0.01 0.13 0.11230 0.00092 0.05364 0.00045 307.5 2.7 1 in rim oz

7.1 123 0.30 0.99 0.10213 0.00094 0.06115 0.00097 314.3 2.8 1 rim oz

8.1 51 0.44 0.24 0.29272 0.00279 0.06797 0.00071 805.4 7.4 1 rim oz

E.1.3 Ios Orthogneiss (Z2405, 89640)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


1.1 433 0.53 0.75 0.15565 0.00202 0.06106 0.00245 403.8 5.4 1 rim oz

2.1 134 0.44 2.17 0.10027 0.00529 0.07312 0.00758 416.9 22.4 1 term oz

3.1 96 0.43 6.21 0.13255 0.00312 0.10455 0.01123 351.3 8.0 1 term oz

3.2 54 0.37 3.39 0.12645 0.00194 0.08073 0.00229 323.0 4.9 1 core oz

4.1 382 0.31 0.28 0.13811 0.00140 0.05564 0.00068 341.3 3.6 1 centre oz

5.1 89 0.47 4.46 0.13485 0.00205 0.08961 0.00669 326.5 5.0 1 term oz

6.1 47 0.63 2.60 0.13080 0.00261 0.07438 0.00249 327.8 6.4 1 centre oz

7.1 85 0.42 2.43 0.13481 0.00184 0.07279 0.00389 319.8 4.5 1 edge oz

8.1 39 0.50 5.73 0.12967 0.00266 0.09988 0.00846 317.9 6.4 1 centre oz

9.1 438 0.22 0.61 0.13696 0.00157 0.05806 0.00080 328.0 4.0 1 centre oz

10.1 120 0.30 3.60 0.13544 0.00282 0.08260 0.01044 329.4 6.8 1 term oz

11.1 33 0.53 5.88 0.13524 0.00456 0.10097 0.00637 310.5 10.3 1 edge oz

12.1 103 0.33 2.12 0.13132 0.00197 0.06986 0.00166 300.6 4.7 1 term oz

Appendix E 275

E.1.4 IO9607 Ios Leucogneiss (Z2665, 97761)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1111σσσσ Age

(Ma)


1.1 67 0.28 0.11 0.22064 0.00861 0.05960 0.00483 545.0 20.4 1 core oxc

2.1 404 0.36 -0.07 0.12918 0.00198 0.05227 0.00133 317.6 4.8 1 rim oz

3.1 49 0.49 0.28 0.11216 0.00436 0.05480 0.00348 296.2 11.3 1 rim oz

4.1 47 1.64 -1.31 1.05836 0.01417 0.15897 0.00508 2436.7 77.6 2 core oxc

5.1 120 0.89 1.27 0.17400 0.00356 0.06595 0.00253 411.2 8.2 1 centre oz

6.1 177 0.25 0.06 0.17485 0.00310 0.05569 0.00214 410.9 7.1 1 centre oz

7.1 214 0.04 0.08 0.12342 0.00179 0.05305 0.00113 295.6 4.2 1 term oz

8.1 112 0.23 4.72 0.48116 0.02136 0.11555 0.00189 1087.6 44.5 1 core oxc

8.2 446 0.02 -0.02 0.10856 0.00288 0.05318 0.00135 338.9 9.1 1 term oz

8.3 244 0.69 0.18 0.53498 0.01573 0.08161 0.00052 1258.9 17.8 2 core oxc

9.1 652 0.01 -0.21 0.14032 0.01640 0.05383 0.00085 428.8 48.5 1 term oz

10.1 145 0.37 0.15 0.13292 0.00463 0.05388 0.00086 308.6 10.5 1 term oz

11.1 320 0.08 5.70 0.13693 0.01132 0.10714 0.00242 576.1 46.3 1 core iz

E.1.5 IO9606 Ios Garnet Mica Schist (Z2665, 97762)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 210 0.82 -0.13 0.27108 0.00308 0.05899 0.00219 595.7 6.5 1 core oz

2.1 159 1.27 0.17 0.28339 0.00361 0.06028 0.00065 550.3 7.1 1 rim oz

3.1 322 0.22 0.37 0.14943 0.00698 0.05839 0.00244 413.0 18.8 1 edge oz

4.1 176 0.13 -0.28 0.13870 0.00977 0.05319 0.00654 428.4 29.3 1 term oz

5.1 217 0.05 0.87 0.96743 0.01063 0.11456 0.00308 1859.0 54.9 2 core oxc

6.1 356 0.81 0.09 0.24639 0.00514 0.05965 0.00130 553.6 11.1 1 edge oz

7.1 222 0.30 0.78 0.23304 0.00821 0.06390 0.00468 493.1 16.7 1 term oz

8.1 251 0.05 1.69 0.32838 0.01566 0.07723 0.00288 695.9 31.5 1 term oz

9.1 179 0.08 1.75 0.30410 0.00667 0.07635 0.00163 657.1 13.8 1 core iz

10.1 55 0.35 0.65 0.30136 0.01368 0.07320 0.00119 853.4 36.4 1 edge iz

11.1 398 0.04 7.43 0.23563 0.01300 0.12288 0.00219 627.7 33.1 1 edge oo

E.1.6 IO9609 Ios Garnet Mica Schist (Z2665, 97763)


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


1.1 276 0.05 0.96 0.12058 0.01106 0.06148 0.00089 341.6 30.5 1 term oz

1.2 64 0.49 0.08 0.39149 0.00619 0.07278 0.00063 977.9 14.4 1 rim oz

2.1 124 0.01 0.87 0.05750 0.00615 0.05906 0.00301 269.9 28.6 1 term oz

2.2 201 0.46 -0.11 0.40237 0.00806 0.07303 0.00112 1027.8 19.1 1 rim oz

3.1 225 0.15 0.19 0.17923 0.01143 0.05889 0.00057 497.7 30.6 1 term oz

3.2 102 0.82 0.01 0.25293 0.00240 0.06056 0.00063 613.4 5.5 1 core oz

4.1 213 0.02 0.41 0.11340 0.01037 0.05685 0.00165 342.9 30.6 1 term oz

5.1 40 2.05 0.27 0.25439 0.00349 0.06199 0.00117 586.0 7.7 1 core iz

6.1 248 0.08 0.15 0.15781 0.00423 0.05600 0.00100 396.3 10.3 1 term oz

7.1 23 0.61 0.71 0.22087 0.00478 0.06565 0.00193 586.7 12.2 1 centre oz

E.1.7 PA9606 Paros Orthogneiss (Z2644, 97764)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 239 0.31 0.33 0.21029 0.00503 0.05507 0.00089 289.7 7.2 1 rim oz

2.1 341 0.05 0.24 0.20823 0.00099 0.05460 0.00088 299.0 2.6 1 rim oz

3.1 525 0.31 0.09 0.11412 0.00287 0.05262 0.00106 267.6 6.7 1 rim oz

4.1 285 0.15 0.30 0.08944 0.00249 0.05357 0.00219 236.7 6.7 1 rim oz

5.1 386 0.21 0.11 0.15189 0.00299 0.05310 0.00064 280.3 5.4 1 rim oz

6.1 162 0.32 0.40 0.16486 0.00130 0.05588 0.00232 298.5 2.3 1 rim oz

7.1 84 0.27 0.56 0.17136 0.00389 0.05717 0.00219 297.7 6.6 1 rim oz

8.1 96 0.10 0.55 0.19008 0.00277 0.05732 0.00215 306.3 4.6 1 rim oz

9.1 440 0.24 0.12 0.17696 0.00163 0.05354 0.00090 298.2 2.9 1 term oz

10.1 444 0.20 0.09 0.16930 0.00318 0.05324 0.00092 294.2 5.5 1 rim oz

11.1 181 0.12 0.21 0.17527 0.00241 0.05426 0.00179 295.9 4.1 1 rim oz

12.1 79 0.26 0.62 0.26170 0.00515 0.06018 0.00130 400.7 7.9 1 core oxc

13.1 425 0.39 0.09 0.20918 0.00200 0.05348 0.00099 303.6 3.6 1 rim oz

14.1 417 0.01 0.29 0.16919 0.00280 0.05495 0.00118 296.9 4.9 1 term oz

15.1 153 0.20 0.34 0.28667 0.00726 0.05897 0.00166 439.9 11.0 1 core oxc

16.1 393 0.06 0.02 0.19866 0.00259 0.05284 0.00117 300.0 4.2 1 rim oz

17.1 366 0.10 0.07 0.17502 0.00157 0.05276 0.00090 282.2 2.8 1 rim oz

18.1 542 0.02 0.00 0.16010 0.00218 0.05294 0.00054 313.2 4.2 1 rim oz

19.1 25 0.40 1.71 0.16938 0.00647 0.06642 0.00357 291.9 10.9 1 core oz

Appendix E 277

E.1.8 PA9601 Paros Orthogneiss (Z2665, 97765)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


2.1 90 0.44 0.56 0.13517 0.00182 0.05777 0.00269 333.1 4.5 1 term oz

3.1 151 0.53 0.55 0.14164 0.00223 0.05780 0.00134 337.6 5.3 1 centre oz

4.1 227 0.21 0.94 0.18747 0.00575 0.06373 0.00116 448.0 13.4 1 core oxc

5.1 643 0.12 0.08 0.12813 0.00255 0.05301 0.00131 299.8 5.9 1 term oz

6.1 790 0.09 0.33 0.09770 0.00368 0.05477 0.00092 284.3 10.7 1 term oz

7.1 213 0.15 0.20 0.14381 0.00152 0.05474 0.00165 330.8 3.5 1 core oxc

7.2 242 0.15 0.30 0.12916 0.00488 0.05492 0.00102 302.3 11.2 1 term oz

8.1 140 0.40 0.32 0.14216 0.00316 0.05562 0.00155 326.6 7.1 1 core oz

E.1.9 SK9601 Sikinos Orthogneiss (Z2633, 97766)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 106 0.30 0.17 0.13112 0.00144 0.05401 0.00127 311.2 3.6 1 rim oz

2.1 381 0.10 0.10 0.12294 0.00341 0.05354 0.00057 313.8 8.5 1 term oz

2.2 739 0.11 0.04 0.11816 0.00599 0.05404 0.00029 357.2 17.6 1 term oz

3.1 43 0.25 5.26 0.07528 0.00399 0.09785 0.00883 387.5 20.9 1 term oz

5.1 346 0.13 0.40 0.11406 0.00382 0.05675 0.00053 342.4 11.2 1 term oz

5.2 78 0.22 0.14 0.11500 0.00173 0.05396 0.00148 315.0 4.6 1 core oz

6.1 1147 0.08 0.52 0.13427 0.00308 0.05739 0.00083 330.7 7.4 1 rim oz

7.1 194 0.32 2.74 0.15159 0.00131 0.07573 0.00188 332.6 3.2 1 term oz

8.1 172 0.18 0.07 0.13271 0.00116 0.05370 0.00089 329.1 2.9 1 term oz

9.1 246 0.13 0.14 0.12348 0.00067 0.05436 0.00052 332.6 1.8 1 term oz

10.1 314 0.16 0.04 0.12125 0.00202 0.05283 0.00039 332.9 5.4 1 term oz

11.1 197 0.18 0.08 0.12588 0.00248 0.05352 0.00063 320.3 6.2 1 rim oz

12.1 527 0.11 0.01 0.12045 0.00466 0.05406 0.00074 370.5 14.0 1 centre oz

13.1 504 0.08 0.04 0.11336 0.00357 0.05437 0.00048 368.5 11.4 1 term oz

14.1 105 0.35 0.68 0.14134 0.00196 0.05839 0.00215 314.5 4.5 1 rim iz

15.1 159 0.86 0.19 0.13349 0.00117 0.05418 0.00084 309.6 2.9 1 in rim oz

16.1 132 0.13 0.41 0.12413 0.00095 0.05597 0.00068 308.2 2.4 1 rim oz

17.1 193 0.03 0.51 0.37693 0.02725 0.07215 0.00293 858.6 58.2 1 core oxc

18.1 222 0.13 0.28 0.12087 0.00133 0.05470 0.00133 299.1 3.3 1 rim oz

17.1 48 0.45 13.58 0.18486 0.00695 0.16407 0.03393 289.3 10.7 1 core oz

18.1 249 0.17 5.25 0.21888 0.00595 0.09649 0.01856 338.2 9.3 1 rim oz

19.1 42 0.40 5.91 0.18068 0.00570 0.10085 0.00697 287.3 9.1 1 in rim oz


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


20.1 170 0.18 2.89 0.18511 0.00339 0.07658 0.01818 316.2 5.9 1 rim oz

21.1 233 0.12 1.62 0.18149 0.00442 0.06595 0.00723 309.6 7.6 1 rim oz

22.1 159 0.19 3.15 0.17623 0.00315 0.07870 0.00709 315.2 5.7 1 rim oz

23.1 841 0.06 1.23 0.19668 0.00394 0.06301 0.00159 322.5 6.7 1 term oz

24.1 51 0.30 9.19 0.17167 0.00485 0.12787 0.02339 288.3 8.1 1 core oz

25.1 403 0.20 1.71 0.18184 0.00338 0.06663 0.00650 306.1 5.9 1 term oz

26.1 43 0.30 9.08 0.18929 0.00752 0.12720 0.02288 299.8 11.8 1 rim oz

E.1.10 NX9314 Naxos Layered Acid Gneiss (Z1889, 97767)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 1362 0.28 0.32 0.09873 0.00180 0.05566 0.00115 324.1 5.8 1 in rim oz

2.1 282 0.62 0.88 0.09761 0.00252 0.06023 0.00121 322.9 8.1 1 centre oz

3.1 79 0.35 5.87 0.09412 0.00393 0.10226 0.00576 352.1 14.4 1 edge oz

4.1 118 0.37 1.04 0.09462 0.00203 0.06149 0.00124 319.4 6.7 1 centre oz

5.1 424 0.43 0.38 0.08694 0.00280 0.05592 0.00059 314.0 9.9 1 term oz

6.1 315 0.45 0.23 0.08922 0.00370 0.05468 0.00207 314.2 12.7 1 centre oz

7.1 253 0.48 0.00 0.09146 0.00357 0.05288 0.00065 319.1 12.1 1 rim oz

8.1 130 0.26 0.65 0.07756 0.00281 0.05801 0.00093 306.4 10.9 1 edge oz

9.1 204 0.28 0.18 0.10658 0.00939 0.05499 0.00382 343.8 29.5 1 rim oz

10.1 240 0.38 1.02 0.08086 0.00177 0.06027 0.00113 271.9 5.8 1 rim oz

11.1 115 0.33 1.63 0.08784 0.00256 0.06636 0.00159 317.7 9.1 1 rim oz

12.1 180 0.52 1.27 0.08758 0.00768 0.06305 0.00285 302.3 25.9 1 centre oz

13.1 583 0.32 0.24 0.10772 0.00205 0.05492 0.00084 320.7 6.0 1 in rim oz

14.1 137 0.52 1.24 0.09626 0.00148 0.06295 0.00156 311.6 4.7 1 centre oz

15.1 215 0.57 0.81 0.08886 0.00396 0.05929 0.00105 306.8 13.4 1 centre oz

16.1 268 0.41 2.54 0.10412 0.00077 0.07397 0.00590 319.2 2.4 1 rim oz

17.1 577 0.42 0.24 0.11444 0.01278 0.05519 0.00421 333.2 36.3 1 term oz

18.1 109 0.43 1.20 0.08546 0.00262 0.06261 0.00147 311.5 9.4 1 centre oz

19.1 122 0.54 1.11 0.08615 0.00237 0.06233 0.00106 328.1 8.8 1 rim oz

20.1 771 0.61 0.42 0.09669 0.00248 0.05635 0.00171 319.2 8.0 1 rim oz

21.1 122 0.47 1.10 0.08991 0.00165 0.06209 0.00096 323.9 5.8 1 term oz

22.1 351 0.19 0.76 0.15977 0.00870 0.06385 0.00280 507.6 26.6 1 rim ixc

22.2 278 0.03 0.81 0.09133 0.00137 0.05971 0.00115 323.1 4.8 1 core oz

Appendix E 279

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


23.1 563 0.11 0.57 0.13670 0.00781 0.06303 0.00089 532.9 29.3 1 core ixc

24.1 153 0.53 0.74 0.09328 0.00372 0.05907 0.00127 320.3 12.4 1 in rim oz

25.1 182 0.34 0.36 0.10231 0.00140 0.05633 0.00074 336.9 4.5 1 edge iz

26.1 144 0.46 0.52 0.10400 0.00293 0.05825 0.00100 361.3 9.9 1 rim oz

E.1.11 NX9485 Naxos Layered Acid Gneiss (Z2645, 97768)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 393 0.53 0.04 0.13430 0.00109 0.05296 0.00100 309.4 2.5 1 rim oz

2.1 255 0.16 0.09 0.13115 0.00184 0.05341 0.00134 313.0 4.3 1 rim oz

3.1 313 0.47 0.16 0.11594 0.00147 0.05355 0.00141 294.1 3.7 1 core oxc

4.1 260 0.45 0.24 0.14365 0.00151 0.05495 0.00156 324.3 3.4 1 term oz

5.1 102 0.41 0.02 0.14883 0.00841 0.05246 0.00429 296.6 16.4 1 term oz

6.1 217 0.09 0.22 0.11922 0.00168 0.05383 0.00141 283.3 3.9 1 in rim oz

7.1 201 0.01 5.01 0.00575 0.00033 0.08793 0.01797 15.0 0.9 1 term oz

8.1 216 0.33 0.25 0.12223 0.00188 0.05447 0.00151 301.0 4.5 1 core oxc

9.1 185 0.41 0.22 0.11485 0.00151 0.05417 0.00157 300.0 3.9 1 term oz

10.1 203 0.09 0.11 0.25192 0.00570 0.05973 0.00105 555.1 12.1 1 core oxc

10.1 740 0.02 0.23 0.11279 0.00379 0.05335 0.00125 260.4 8.6 1 term oz

10.2 125 0.63 0.29 0.10628 0.00445 0.04970 0.00853 286.8 11.8 1 core oz

11.1 499 0.21 0.15 0.15057 0.00526 0.05338 0.00143 289.7 10.0 1 term oz

11.2 93 0.53 0.08 0.14890 0.00249 0.05323 0.00283 306.8 5.2 1 core oz

12.1 309 0.29 0.35 0.12338 0.00320 0.05513 0.00156 293.4 7.5 1 term oz

E.1.12 NX9315 Naxos Leucogneiss (Z2264, 97769)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 205 0.03 0.42 0.05473 0.00245 0.05277 0.00200 160.8 7.1 1 rim oz

1.2 213 0.02 2.10 0.01169 0.00042 0.06367 0.00391 28.1 1.0 1 term uol

2.1 282 0.05 0.53 0.05160 0.00078 0.05286 0.00130 122.8 1.8 1 edge oz

3.1 293 0.01 0.38 0.00662 0.00011 0.04953 0.00185 19.3 0.3 1 term uol

4.1 137 0.25 0.42 0.06519 0.00116 0.05301 0.00122 172.7 3.0 1 edge oz

5.1 196 0.08 0.77 0.07072 0.00245 0.05590 0.00164 173.1 5.9 1 term oz


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


6.1 247 0.01 1.20 0.00688 0.00020 0.05609 0.00259 16.9 0.5 1 term uol

6.2 206 0.01 1.33 0.00687 0.00026 0.05723 0.00291 18.0 0.7 1 term uol

7.1 238 0.09 0.40 0.07654 0.00115 0.05364 0.00105 205.4 3.1 1 term oz

8.1 169 0.02 0.88 0.01453 0.00065 0.05401 0.00242 39.3 1.8 1 rim uol

9.1 407 0.01 1.44 0.00717 0.00009 0.05814 0.00287 19.4 0.2 1 term uol

9.2 369 0.01 0.07 0.00686 0.00020 0.04591 0.00260 19.8 0.6 1 term uol

10.1 172 0.15 1.44 0.04587 0.00209 0.05977 0.00295 102.7 4.6 1 term uol

11.1 322 0.03 1.25 0.01510 0.00068 0.05716 0.00257 49.2 2.2 1 rim mzo

12.1 407 0.04 1.90 0.01080 0.00066 0.06209 0.00394 30.4 1.8 1 rim uol

12.2 433 0.14 0.47 0.06693 0.00127 0.05331 0.00107 172.1 3.2 1 centre oz

13.1 593 0.01 0.76 0.01848 0.00048 0.05300 0.00161 41.6 1.1 1 rim oo

14.1 382 0.03 1.99 0.01122 0.00027 0.06274 0.00307 27.7 0.7 1 rim mix

15.1 664 0.01 0.75 0.00695 0.00009 0.05250 0.00202 17.9 0.2 1 rim uol

15.2 484 0.01 1.06 0.00757 0.00014 0.05500 0.00298 18.6 0.3 1 rim uol

16.1 703 0.01 1.58 0.00722 0.00017 0.05918 0.00520 17.4 0.4 1 rim oo

17.1 383 0.02 1.26 0.02271 0.00106 0.05740 0.00236 56.3 2.6 1 rim mzo

18.1 543 0.02 2.11 0.01112 0.00029 0.06373 0.00223 30.2 0.8 1 rim mix

19.1 449 0.14 0.79 0.04470 0.00109 0.05474 0.00172 116.2 2.8 1 rim oz

20.1 50 0.11 4.04 0.02947 0.00143 0.08038 0.00761 78.3 3.8 1 rim iz

21.1 315 0.20 0.91 0.04874 0.00063 0.05583 0.00226 121.6 1.6 1 rim uol

22.1 472 0.04 0.66 0.06837 0.00199 0.05492 0.00210 177.2 5.1 1 rim oz

23.1 215 0.08 3.66 0.06068 0.00209 0.07832 0.00463 133.2 4.6 1 rim oz

24.1 447 0.03 2.28 0.01847 0.00050 0.06538 0.00746 42.7 1.2 1 term oz

25.1 382 0.04 2.78 0.01751 0.00060 0.06952 0.00664 47.7 1.6 1 rim uol

1.1a 756 0.02 1.13 0.02098 0.00053 0.05631 0.00181 50.4 1.3 1 rim mzo

2.1a 564 0.07 0.56 0.03617 0.00053 0.05253 0.00155 98.6 1.4 1 rim rz

2.2a 845 0.03 0.73 0.04180 0.00097 0.05402 0.00075 104.0 2.4 1 rim rz

3.1a 574 0.20 0.81 0.04344 0.00325 0.05474 0.00128 106.2 7.9 1 rim uol

4.1a 1058 0.01 0.51 0.00625 0.00008 0.05055 0.00138 17.2 0.2 1 rim uol

5.1a 713 0.08 0.78 0.03813 0.00116 0.05501 0.00154 129.2 3.9 1 rim uol

Appendix E 281


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 85 0.01 3.29 0.00700 0.00048 0.07313 0.00681 19.0 1.3 1 rim uol

1.2 129 0.01 6.43 0.00528 0.00022 0.09857 0.01278 19.6 0.8 1 rim uol

2.1 95 0.02 3.45 0.00822 0.00049 0.07461 0.00672 27.9 1.7 1 rim uol

3.1 209 0.22 0.28 0.09764 0.00101 0.05366 0.00124 259.7 2.6 1 rim oo

4.1 1108 0.08 2.01 0.02015 0.00029 0.06342 0.00313 54.1 0.8 1 rim rz

5.1 117 0.03 2.18 0.01426 0.00025 0.06450 0.00272 39.5 0.7 1 rim uol

6.1 167 0.07 4.77 0.02906 0.00127 0.08620 0.00856 77.5 3.4 1 rim uol

7.1 86 0.14 1.98 0.02897 0.00063 0.06403 0.00233 100.3 2.2 1 rim mix

8.1 290 0.01 1.33 0.00688 0.00016 0.05720 0.00344 17.9 0.4 1 rim uol

8.2 305 0.00 2.40 0.00578 0.00019 0.06585 0.00262 17.0 0.6 1 rim uol

8.3 262 0.01 2.15 0.00576 0.00017 0.06386 0.00278 17.4 0.5 1 rim uol

8.4 311 0.00 1.64 0.00574 0.00012 0.05971 0.00350 17.6 0.4 1 rim uol

9.1 115 0.02 1.72 0.00778 0.00018 0.06045 0.00356 23.0 0.5 1 rim uol

10.1 134 0.06 1.31 0.01339 0.00037 0.05751 0.00348 41.4 1.2 1 rim rz

11.1 723 0.21 0.45 0.08992 0.00110 0.05449 0.00059 233.7 2.8 1 rim lm

13.1 273 0.01 1.69 0.00636 0.00013 0.06012 0.00303 18.4 0.4 1 rim uol

13.2 220 0.01 2.30 0.00587 0.00013 0.06502 0.00451 17.6 0.4 1 rim uol

14.1 736 0.28 0.30 0.10184 0.00059 0.05396 0.00048 265.2 1.5 1 edge oz

15.1 352 0.14 0.33 0.05800 0.00338 0.05236 0.00180 180.8 10.4 1 edge oz

16.1 128 0.02 2.33 0.01057 0.00036 0.06556 0.00380 30.8 1.1 1 rim uol

16.2 279 0.01 2.98 0.00556 0.00022 0.07058 0.00468 18.4 0.7 1 rim uol

17.1 308 0.14 0.57 0.08785 0.00068 0.05535 0.00092 227.9 1.7 1 term oz

17.2 451 0.26 -0.03 0.13137 0.00133 0.05251 0.00089 319.5 3.2 1 in rim oz

18.1 297 0.02 2.77 0.00647 0.00014 0.06884 0.00365 17.8 0.4 1 rim uol

18.2 290 0.02 2.32 0.01483 0.00093 0.06571 0.00447 41.9 2.6 1 rim uol

19.1 210 0.01 2.54 0.00697 0.00018 0.06703 0.00365 17.4 0.4 1 term uol

19.2 78 0.08 3.12 0.05146 0.00273 0.07354 0.00592 100.8 5.3 1 rim oo

20.1 145 0.02 1.52 0.00562 0.00032 0.05882 0.00755 16.8 1.0 1 rim uol

20.2 320 0.40 0.20 0.11053 0.00438 0.05368 0.00151 281.6 11.1 1 core oxc

21.1 626 0.14 0.69 0.06908 0.00411 0.05521 0.00111 170.2 10.0 1 rim lm

22.1 199 0.02 4.93 0.00634 0.00035 0.08672 0.00840 18.1 1.0 1 rim oo

23.1 141 0.06 12.42 0.02324 0.00120 0.14846 0.02007 44.5 2.3 1 rim uol

23.2 92 0.01 17.31 0.00644 0.00095 0.18798 0.04694 15.3 2.3 1 rim uol

23.3 389 0.30 0.43 0.15200 0.00347 0.05588 0.00147 294.5 6.6 1 rim lm


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


24.1 689 0.24 0.20 0.14977 0.00121 0.05438 0.00098 311.8 2.5 1 rim uol

25.1 620 0.26 0.20 0.17287 0.00196 0.05446 0.00102 315.2 3.5 1 rim oz

26.1 126 0.19 1.60 0.08610 0.00293 0.06272 0.00585 175.4 5.9 1 rim oo

27.1 359 0.40 0.41 0.13006 0.00378 0.05604 0.00286 306.6 8.8 1 rim oo

28.1 136 0.01 15.37 0.00734 0.00045 0.17208 0.06113 15.9 1.0 1 rim uol

29.1 165 0.02 11.16 0.00843 0.00051 0.13764 0.02229 16.1 1.0 1 rim uol

30.1 456 0.08 0.19 0.15651 0.00219 0.05409 0.00172 301.9 4.1 1 rim oz


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 99 0.40 0.30 0.11357 0.00244 0.05512 0.00159 314.2 6.6 1 term oz

2.1 317 0.38 0.05 0.11323 0.00153 0.05320 0.00073 319.0 4.2 1 rim oz

2.2 68 0.37 0.73 0.10781 0.00137 0.05800 0.00183 288.1 3.6 1 rim oz

3.1 91 0.47 0.61 0.11387 0.00086 0.05784 0.00124 322.6 2.4 1 term oz

4.1 134 0.15 0.28 0.07783 0.00226 0.05302 0.00134 228.3 6.5 1 term uol

5.1 49 0.48 0.78 0.11687 0.00247 0.05911 0.00161 320.6 6.6 1 rim oz

5.2 84 0.46 0.50 0.12608 0.00275 0.05713 0.00145 331.7 7.1 1 centre oz

7.1 240 0.51 0.22 0.11175 0.00204 0.05440 0.00146 313.4 5.6 1 rim oz

8.1 99 0.04 3.15 0.01354 0.00063 0.07255 0.00803 36.0 1.7 1 term uol

9.1 309 0.36 0.07 0.09644 0.00248 0.05171 0.00171 298.4 7.5 1 term oz

10.1 103 0.01 8.26 0.00631 0.00033 0.11400 0.01871 18.6 1.0 1 rim uol

11.1 185 0.02 1.40 0.00489 0.00022 0.03493 0.00630 17.4 0.8 1 rim uol

12.1 89 0.14 4.97 0.05238 0.00120 0.00834 0.00088 144.4 3.3 1 rim uol

13.1 230 0.24 5.74 0.07178 0.00109 0.00361 0.00032 212.9 3.2 1 term oz

14.1 1191 0.26 5.92 0.09207 0.00273 0.00279 0.00066 241.8 7.1 1 rim oo

15.1 171 0.43 0.46 0.11396 0.00286 0.05598 0.00146 297.0 7.3 1 rim oz

16.1 136 0.03 1.70 0.03907 0.00154 0.06218 0.00247 114.6 4.5 1 rim uol

17.1 198 0.24 0.37 0.10919 0.00364 0.05472 0.00261 273.1 8.9 1 rim uol

18.1 1027 0.39 0.01 0.13717 0.00365 0.05331 0.00160 339.0 8.8 1 rim oz

19.1 975 0.34 0.28 0.12446 0.00289 0.05475 0.00119 306.5 7.0 1 rim oz

20.1 1580 0.06 0.48 0.04794 0.00151 0.05259 0.00072 130.6 4.1 1 rim mzo

21.1 1507 0.05 17.56 0.04909 0.00265 0.19205 0.02266 132.4 7.1 1 rim mzo

22.1 1278 0.16 0.20 0.08149 0.00096 0.05289 0.00066 252.4 3.0 1 rim mix

Appendix E 283

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


23.1 313 0.27 0.23 0.15457 0.00591 0.05458 0.00114 319.8 12.0 1 rim oz

24.1 511 0.14 0.60 0.08645 0.00526 0.05569 0.00163 235.3 14.1 1 rim oz

25.1 94 0.42 0.52 0.12269 0.00524 0.05637 0.00163 294.9 12.3 1 term oz

26.1 129 0.02 20.08 0.00695 0.00072 0.20952 0.04314 15.4 1.6 1 term uol

26.2 114 0.01 0.66 0.00455 0.00035 0.05166 0.01384 14.9 1.1 1 term uol

26.3 54 0.39 1.53 0.11401 0.00451 0.06413 0.00269 277.2 10.7 1 centre oz

27.1 51 0.16 4.05 0.04441 0.00151 0.08094 0.00562 103.9 3.5 1 rim mix

27.2 72 0.01 10.52 0.00657 0.00054 0.13180 0.02040 14.5 1.2 1 rim uol

28.1 66 0.02 7.35 0.01169 0.00067 0.10624 0.01782 26.7 1.5 1 term uol

29.1 168 0.02 8.35 0.00771 0.00025 0.11420 0.01519 17.1 0.5 1 rim uol

29.2 106 0.03 4.40 0.01183 0.00042 0.08234 0.00706 28.3 1.0 1 rim uol

E.1.15 NX94103 Naxos Migmatite (Z2153, 97772)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 234 0.19 1.17 0.08965 0.00321 0.06077 0.00063 251.8 8.8 1 term oz

2.1 23 0.24 9.72 0.08196 0.00197 0.12855 0.00619 180.2 4.4 1 term oz

3.1 45 0.24 3.88 0.10303 0.00277 0.08253 0.00466 244.6 6.5 1 rim oz

4.1 268 0.02 2.22 0.02779 0.00099 0.06544 0.00190 66.0 2.4 1 term oz

5.1 893 0.09 0.06 0.14248 0.00398 0.05334 0.00026 321.3 9.0 1 core oxc

6.1 111 0.04 0.82 0.14235 0.00693 0.06018 0.00200 351.7 16.7 1 rim oz

7.1 334 0.09 2.92 0.03334 0.00122 0.07188 0.00176 105.3 3.8 1 term oz

8.1 1202 0.03 0.21 0.09908 0.00391 0.05364 0.00031 280.8 10.9 1 rim oz

9.1 80 0.29 1.78 0.13398 0.00182 0.06686 0.00117 306.4 4.4 1 rim oz

10.1 163 0.07 2.35 0.04643 0.00135 0.06791 0.00103 137.4 4.0 1 rim uol

11.1 513 0.16 0.53 0.11075 0.00353 0.05656 0.00049 297.4 9.3 1 rim uol

12.1 667 0.02 1.16 0.06340 0.00390 0.05963 0.00054 205.0 12.4 1 rim oz

13.1 182 0.15 0.86 0.12583 0.00182 0.05942 0.00087 306.6 4.5 1 term oz

14.1 354 0.03 2.48 0.04544 0.00219 0.06892 0.00102 137.1 6.5 1 term uol

15.1 452 0.02 5.14 0.01232 0.00029 0.08842 0.00383 27.8 0.7 1 rim uol

16.1 559 0.11 0.59 0.10396 0.00163 0.05674 0.00041 281.9 4.3 1 rim uol

17.1 290 0.02 9.81 0.00935 0.00023 0.12656 0.00260 31.8 0.8 1 rim uol

18.1 148 0.07 2.47 0.05349 0.00159 0.06862 0.00175 123.7 3.7 1 rim uol

19.1 203 0.13 1.10 0.11844 0.00173 0.06119 0.00131 296.2 4.3 1 term oz

20.1 349 0.07 1.66 0.04980 0.00195 0.06422 0.00141 229.2 9.2 1 term oz


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


21.1 298 0.13 1.17 0.07347 0.00209 0.06052 0.00092 243.4 6.9 1 rim oz

22.1 161 0.33 1.57 0.11833 0.00460 0.06498 0.00086 297.9 11.4 1 rim oz

23.1 201 0.23 0.97 0.09244 0.00094 0.05951 0.00064 269.2 2.7 1 rim oz

24.1 213 0.07 2.42 0.05593 0.00105 0.06824 0.00129 126.9 2.5 1 rim uol

25.1 132 0.03 10.19 0.01375 0.00017 0.12959 0.00566 29.0 0.4 1 rim uol

26.1 330 0.06 1.92 0.02764 0.00097 0.06292 0.00463 63.2 2.2 1 rim uol


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 361 0.01 0.77 0.11847 0.00225 0.05812 0.00225 259.7 4.8 1 rim oz

2.1 456 0.01 6.14 0.03987 0.00205 0.10047 0.01789 90.8 4.6 1 rim mix

2.2 76 0.30 1.57 0.18384 0.00534 0.06839 0.00259 406.6 11.4 1 core ixc

2.3 827 0.00 0.29 0.00693 0.00015 0.04877 0.00198 14.8 0.3 1 term uol

3.1 127 0.19 0.39 0.14381 0.00313 0.05610 0.00194 312.7 6.7 1 core oz

3.2 466 0.06 0.45 0.10169 0.00422 0.05662 0.00165 315.1 13.0 1 rim oz

4.1 611 0.01 2.97 0.01953 0.00155 0.07256 0.00910 51.7 4.1 1 rim uol

5.1 393 0.00 3.18 0.04829 0.00225 0.07622 0.01082 143.5 6.7 1 term oz

6.1 609 0.03 2.77 0.02394 0.00042 0.07087 0.00434 53.9 1.0 1 rim uol

6.2 54 0.58 0.88 0.29246 0.00456 0.06988 0.00166 677.2 10.0 1 core oxc

7.2 583 0.06 0.14 0.16505 0.00122 0.05588 0.00094 392.3 2.8 1 core oz

8.1 344 0.02 0.57 0.10578 0.00175 0.05613 0.00116 248.8 4.0 1 term oz

9.1 457 0.31 1.56 0.07323 0.00175 0.06306 0.00528 178.5 4.2 1 term oz

10.1 348 0.01 0.12 0.12588 0.00188 0.05330 0.00159 292.4 4.3 1 term oz

11.1 889 0.00 1.34 0.00646 0.00022 0.05786 0.00329 15.4 0.5 1 term uol

11.2 75 0.95 0.27 0.43862 0.00525 0.07761 0.00132 1058.6 11.9 1 core oxc

12.1 289 0.04 0.41 0.11060 0.00168 0.05610 0.00129 304.0 4.7 1 term oz

13.1 611 0.00 2.26 0.00865 0.00085 0.06596 0.00905 26.1 2.6 1 term uol

13.2 151 0.62 6.20 0.51504 0.00514 0.13252 0.00140 1871.2 30.5 2 core oxc

13.3 343 0.00 1.60 0.01768 0.00155 0.06082 0.00455 60.0 5.3 1 term uol

14.1 572 0.01 1.48 0.06103 0.00205 0.06153 0.00395 137.1 4.6 1 term oo

14.2 185 0.26 0.43 0.15287 0.00195 0.05827 0.00132 388.7 4.9 1 core oxc

15.2 61 0.43 0.15 0.11671 0.00263 0.05403 0.00301 313.1 7.0 1 core oxc

15.3 609 0.01 0.05 0.15533 0.00186 0.05321 0.00079 314.6 4.0 1 term uol

Appendix E 285

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


16.1 198 0.08 0.26 0.08249 0.00428 0.05406 0.00118 277.1 14.2 1 term oo

17.1 269 0.01 0.15 0.09472 0.00220 0.05322 0.00069 386.9 9.9 1 term oz

18.1 684 0.00 0.40 0.01254 0.00029 0.05013 0.00105 34.3 0.8 1 term uol

19.1 715 0.07 0.37 0.05258 0.00666 0.05375 0.00099 220.9 27.6 1 edge mzo

19.2 2466 0.00 0.06 0.12034 0.00415 0.05264 0.00070 287.8 9.7 1 rim mzo

20.1 1211 0.00 0.29 0.09991 0.00396 0.05288 0.00065 210.8 8.3 1 centre iz

21.1 178 0.05 0.38 0.11693 0.00405 0.05477 0.00075 262.2 8.9 1 centre oz

22.1 1091 0.00 0.28 0.09312 0.00345 0.05300 0.00113 223.3 8.1 1 edge oz

E.1.17 NX9637 Melt Pod Naxos Migmatite (Z2782, 97774)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 633 0.00 0.38 0.00474 0.00022 0.04945 0.00235 18.6 0.9 1 term uol

2.1 984 0.01 1.65 0.00531 0.00017 0.05977 0.00255 18.1 0.6 1 term uol

3.1 335 0.01 2.03 0.00411 0.00053 0.06293 0.01175 20.7 2.7 1 rim uol

4.1 499 0.01 0.04 0.00448 0.00032 0.04668 0.00311 16.7 1.2 1 rim uol

5.1 1186 0.01 0.39 0.00447 0.00044 0.04953 0.00192 16.4 1.6 1 rim uol

6.1 74 0.27 0.55 0.07361 0.00098 0.05509 0.00156 225.1 2.9 1 in rim oxc

5.2 48 0.55 0.11 0.11424 0.00242 0.05395 0.00155 332.5 6.9 1 core oxc

7.2 609 0.74 0.01 0.08900 0.00120 0.05148 0.00052 266.1 3.5 1 core oxc

8.1 1052 0.01 0.66 0.00526 0.00024 0.05169 0.00120 16.7 0.8 1 term uos

9.1 108 0.01 1.35 0.01008 0.00057 0.05737 0.00582 21.0 1.2 1 rim iz

10.1 166 0.06 0.57 0.01995 0.00027 0.05170 0.00172 52.9 0.8 1 in rim oz

11.1 619 0.01 6.22 0.00537 0.00010 0.09687 0.01378 17.3 0.3 1 term uol

12.1 74 0.91 1.61 0.08136 0.00525 0.06556 0.00188 312.5 19.7 1 term oz

13.1 177 0.01 6.27 0.00417 0.00018 0.09733 0.01596 21.9 1.0 1 rim uos

14.1 776 0.00 2.03 0.00539 0.00038 0.06289 0.00246 19.3 1.4 1 term uol

15.1 205 0.01 2.61 0.00578 0.00013 0.06752 0.00429 13.5 0.3 1 rim uol

15.2 24 0.32 0.43 0.09168 0.00302 0.05579 0.00310 298.5 9.6 1 core oxc

16.1 127 0.01 5.62 0.00383 0.00019 0.09201 0.02269 16.5 0.8 1 term uol

17.1 1133 0.01 1.57 0.00675 0.00030 0.05913 0.00382 17.0 0.7 1 term uol

17.2 701 0.01 2.43 0.00353 0.00034 0.06603 0.00578 13.5 1.3 1 edge uol

18.1 97 0.00 4.63 0.00355 0.00016 0.08398 0.01334 17.8 0.8 1 rim uol


E.1.18 NX9451 Naxos Quartzite (Z2156, 97775)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 180 0.40 0.22 0.07295 0.00336 0.05249 0.00245 225.0 10.2 1 rim oz

1.2 124 0.44 1.37 0.08585 0.00261 0.06247 0.00331 249.2 7.5 1 rim oz

2.1 1213 0.26 0.30 0.05665 0.00047 0.05128 0.00145 137.7 1.2 1 rim uol

3.1 556 0.41 -0.33 0.09032 0.00127 0.04827 0.00160 238.8 3.3 1 rim oz

4.1 136 0.35 1.57 0.09019 0.00295 0.06407 0.00420 246.4 7.9 1 rim oz

4.2 84 0.53 2.50 0.08923 0.00274 0.07175 0.00726 249.3 7.5 1 core oxc

6.1 388 0.39 5.00 0.09394 0.00198 0.09195 0.01029 239.6 5.0 1 term oz

6.2 178 0.70 8.71 0.08807 0.00198 0.12237 0.01850 237.6 5.3 1 rim oz

7.1 265 0.42 8.20 0.08707 0.00255 0.11808 0.02241 231.1 6.7 1 rim oz

8.1 168 0.38 7.72 0.08559 0.00215 0.11413 0.01907 231.0 5.7 1 rim oz

9.1 97 0.40 7.79 0.07463 0.00355 0.11442 0.03211 218.6 10.2 1 rim oz

9.2 194 0.46 8.49 0.08097 0.00205 0.12032 0.01340 223.8 5.6 1 core oz

10.1 273 0.43 6.22 0.08072 0.00155 0.10172 0.01658 229.3 4.3 1 edge oz

11.1 398 0.31 4.58 0.08155 0.00141 0.08832 0.00989 229.7 3.9 1 term oz

12.1 276 0.42 4.41 0.07715 0.00179 0.08691 0.01650 230.7 5.3 1 term oz

13.1 184 0.39 9.36 0.08168 0.00258 0.12728 0.02371 215.2 6.7 1 rim oz

14.1 206 0.39 3.46 0.07449 0.00179 0.07890 0.00679 221.6 5.3 1 rim oz

15.1 231 0.50 1.59 0.07589 0.00188 0.06400 0.00816 242.6 6.0 1 rim oz

16.1 218 0.37 3.62 0.08285 0.00184 0.08057 0.01233 237.1 5.2 1 term oz

16.2 190 0.42 6.21 0.08961 0.00207 0.10200 0.01883 244.9 5.6 1 core oz

17.1 266 0.79 5.80 0.08809 0.00191 0.09855 0.01488 240.3 5.1 1 rim oz

17.2 368 0.78 3.36 0.08444 0.00149 0.07827 0.00663 228.4 4.0 1 core oz

18.1 366 0.51 6.46 0.09178 0.00113 0.10409 0.01304 245.6 3.0 1 rim oz

19.1 203 0.59 4.31 0.07825 0.00322 0.08612 0.01234 232.7 9.4 1 core oz

20.1 249 0.43 4.82 0.07940 0.00134 0.09036 0.01515 233.3 3.9 1 term oz

21.1 117 0.38 12.67 0.08923 0.00366 0.15497 0.03538 240.3 9.7 1 rim oz

22.1 271 0.46 4.13 0.07698 0.00126 0.08452 0.01027 226.9 3.7 1 rim oz

23.1 190 0.40 4.95 0.06919 0.00135 0.09108 0.01134 216.4 4.2 1 rim oz

24.1 218 0.43 4.22 0.08604 0.00141 0.08543 0.01006 233.9 3.8 1 rim oz

25.1 408 0.47 3.48 0.08133 0.00107 0.07935 0.01120 231.9 3.0 1 rim oz

26.1 156 0.51 7.92 0.08770 0.00215 0.11578 0.01802 233.6 5.6 1 centre oz

27.1 250 0.39 1.94 0.08145 0.00161 0.06673 0.00245 234.3 4.5 1 term oz

Appendix E 287

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


28.1 185 0.38 9.17 0.08993 0.00151 0.12613 0.03097 235.6 3.9 1 term oz

29.1 292 0.52 6.60 0.08466 0.00153 0.10473 0.01517 223.1 4.0 1 term oz

30.1 279 0.45 5.31 0.08693 0.00160 0.09424 0.01831 227.9 4.1 1 edge oz

31.1 426 0.64 5.71 0.08886 0.00153 0.09769 0.01954 234.3 4.0 1 rim oz

E.1.19 NX9481 Naxos Quartzite (Z2217, 97776)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 211 0.35 0.86 0.12437 0.00281 0.05740 0.00211 210.1 4.9 1 rim oz

2.1 214 0.75 0.69 0.11477 0.00148 0.05643 0.00141 230.4 3.0 1 term oz

3.1 212 0.69 0.78 0.11306 0.00198 0.05699 0.00408 221.3 3.9 1 rim oz

3.2 184 0.62 0.84 0.09337 0.00257 0.05759 0.00183 225.9 6.1 1 term oz

4.1 156 0.57 1.03 0.10494 0.00371 0.05925 0.00246 227.7 7.9 1 rim oz

4.2 371 1.15 0.61 0.10848 0.00107 0.05531 0.00118 208.9 2.2 1 term oz

4.3 301 0.86 0.87 0.11996 0.00181 0.05768 0.00174 216.9 3.4 1 core oxc

5.1 130 0.42 1.79 0.08478 0.00409 0.06481 0.00453 196.6 9.3 1 term oz

5.2 449 0.58 0.48 0.11407 0.00204 0.05469 0.00103 229.3 4.1 1 core oz

5.3 113 0.46 0.72 0.11107 0.00297 0.05658 0.00231 225.4 6.0 1 rim oz

6.1 386 0.70 1.54 0.14617 0.00209 0.06336 0.00333 224.5 3.8 1 rim oz

6.2 397 0.78 0.45 0.10542 0.00319 0.05385 0.00121 200.0 6.0 1 rim oz

7.1 149 0.54 0.39 0.10611 0.00132 0.05400 0.00323 229.2 2.8 1 term oz

8.1 462 1.73 0.24 0.12351 0.00167 0.05287 0.00123 235.7 3.3 1 rim oz

9.1 256 0.28 1.02 0.05287 0.00115 0.05683 0.00325 117.7 2.5 1 rim oz

9.2 273 0.78 0.54 0.11514 0.00113 0.05554 0.00156 243.9 2.4 1 rim oz

10.1 97 0.74 1.04 0.11817 0.00296 0.05944 0.00271 233.1 5.8 1 edge oz

11.1 132 0.44 1.26 0.12880 0.00252 0.06130 0.00307 236.3 4.7 1 term oz

12.1 369 1.00 0.35 0.09601 0.00134 0.05385 0.00143 239.1 3.3 1 rim oz

12.2 277 0.74 0.72 0.10195 0.00119 0.05702 0.00130 242.6 2.8 1 core oz

13.1 163 0.62 0.90 0.11045 0.00218 0.05834 0.00255 238.2 4.6 1 rim oz

14.1 241 0.60 0.56 0.09301 0.00188 0.05500 0.00321 211.3 4.2 1 rim oz

15.1 547 0.03 0.60 0.06263 0.00049 0.05449 0.00246 171.2 1.6 1 rim oz

15.2 117 0.04 1.74 0.14640 0.00307 0.06705 0.00456 315.6 6.5 1 core oxc

15.3 683 0.05 0.89 0.06161 0.00081 0.05635 0.00114 147.7 1.9 1 rim oz

15.4 150 0.10 0.91 0.13182 0.00158 0.06017 0.00213 314.6 3.7 1 core oxc


E.1.20 SY9603 Syros “Vari” Orthogneiss (Z2665, 97777)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 92 0.46 3.23 0.10704 0.00219 0.07712 0.00372 212.3 4.3 1 core oz

2.1 30 0.50 3.94 0.12970 0.00353 0.08361 0.00557 239.9 6.4 1 in rim oz

3.1 52 0.49 1.19 0.10320 0.00568 0.06024 0.00323 211.5 11.5 1 centre iz

4.1 494 0.39 0.09 0.13980 0.00220 0.05204 0.00072 252.8 4.0 1 term sz

5.1 305 0.77 0.26 0.13720 0.00193 0.05336 0.00134 250.1 3.6 1 in rim oz

6.1 177 0.62 0.52 0.10874 0.00108 0.05532 0.00221 239.5 2.3 1 term oz

7.1 89 0.41 1.43 0.11191 0.00207 0.06279 0.00316 235.8 4.3 1 core oz

8.1 89 0.51 0.46 0.11354 0.00170 0.05481 0.00200 240.7 3.5 1 edge oz

9.1 99 0.55 1.31 0.12050 0.00230 0.06184 0.00322 238.9 4.5 1 core oz

10.1 101 0.48 1.42 0.13163 0.00413 0.06287 0.00282 244.6 7.6 1 term oz

11.1 109 0.40 1.21 0.13064 0.00155 0.06114 0.00301 244.1 2.9 1 core oz

12.1 219 0.43 0.27 0.14217 0.00161 0.05359 0.00177 253.5 3.0 1 core oxc

13.1 67 0.59 2.51 0.12386 0.00263 0.07183 0.00420 243.4 5.1 1 core oz

14.1 208 0.40 0.50 0.14324 0.00371 0.05553 0.00281 255.9 6.6 1 edge oz

15.1 173 0.23 0.23 0.12571 0.00611 0.05292 0.00191 240.1 11.5 1 term oz

16.1 34 0.45 4.47 0.11278 0.01121 0.08741 0.00935 211.3 20.7 1 edge rz

17.1 59 0.53 3.14 0.11663 0.00402 0.07707 0.00406 241.7 8.2 1 edge oz

18.1 166 0.44 0.17 0.12387 0.00291 0.05240 0.00148 240.0 5.6 1 rim oz

19.1 73 0.38 1.95 0.12797 0.00364 0.06719 0.00488 242.8 6.8 1 in rim oz

20.1 91 0.56 1.97 0.11143 0.00173 0.06706 0.00392 227.2 3.5 1 centre iz

21.1 103 0.62 1.04 0.11822 0.00242 0.05956 0.00247 236.5 4.8 1 centre oz

22.1 57 0.55 0.88 0.11583 0.00235 0.05839 0.00349 243.4 4.8 1 edge oz

23.1 32 0.16 0.80 0.12580 0.00287 0.05782 0.00240 241.6 5.6 1 term oo

24.1 70 0.56 13.99 0.07266 0.00514 0.16872 0.00843 214.4 15.0 1 rim oz

25.1 432 0.15 0.11 0.09635 0.00568 0.05160 0.00054 222.2 12.9 1 term oz

Appendix E 289

E.1.21 89646 Syros Quartzite (Z2405, 89646)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 321 0.02 0.78 0.24599 0.00467 0.06789 0.00128 650.3 12.2 1 edge lm

2.1 209 0.03 7.19 0.03932 0.00360 0.10851 0.00837 141.9 13.0 1 term oo

3.1 1046 0.28 0.81 0.11229 0.00442 0.06029 0.00225 341.1 13.1 1 term oz

4.1 2991 0.04 4.60 0.03028 0.00076 0.08590 0.00424 83.7 2.1 1 edge IZ

4.2 174 0.42 14.54 0.02931 0.00083 0.16812 0.01423 78.4 2.2 1 centre IZ

5.1 2447 0.01 2.00 0.03575 0.00105 0.06507 0.00255 119.0 3.5 1 term los

6.1 7103 0.07 0.43 0.10591 0.00380 0.05685 0.00041 330.7 11.6 1 edge lm

7.1 1885 0.00 0.24 0.10342 0.00162 0.05472 0.00100 304.9 4.7 1 rim oz

8.1 2186 0.01 0.87 0.03639 0.00097 0.05576 0.00213 121.1 3.2 1 rim oo

9.1 583 0.00 0.92 0.09803 0.00118 0.05993 0.00132 290.5 3.4 1 term oz

10.1 241 0.71 3.19 0.02416 0.00078 0.07358 0.00524 75.9 2.4 1 centre sz

11.1 316 0.02 7.02 0.03220 0.00139 0.10535 0.00844 101.3 4.3 1 rim oz

12.1 177 0.08 1.10 0.14207 0.00141 0.06392 0.00198 408.1 4.2 1 edge oz

13.1 433 0.01 0.59 0.07369 0.00317 0.05709 0.00111 297.5 12.6 1 rim uol

14.1 646 0.04 1.91 0.20969 0.00157 0.07399 0.00066 548.7 4.9 1 rim oz

15.1 347 0.04 3.36 0.09678 0.00234 0.07965 0.00313 298.0 7.0 1 edge oz

16.1 367 0.04 5.23 0.05676 0.00133 0.09325 0.00478 226.2 5.3 1 term oz

17.1 265 0.02 1.72 0.08482 0.00139 0.06598 0.00231 282.8 4.5 1 term oz

18.1 52 0.81 1.85 0.22143 0.00328 0.07551 0.00266 621.4 9.0 1 edge ad

19.1 73 0.42 1.43 0.13728 0.00154 0.06714 0.00273 430.4 4.7 1 edge oz

20.1 472 0.76 0.94 0.08192 0.00180 0.05915 0.00178 261.8 5.6 1 rim oz

21.1 580 0.06 2.15 0.08346 0.00194 0.06958 0.00168 287.6 6.5 1 term oz

22.1 51 0.18 14.65 0.17091 0.00388 0.17433 0.01923 446.0 9.8 1 rim oz

23.1 342 0.06 0.81 0.12255 0.00231 0.06190 0.00123 421.3 7.7 1 rim oz

24.1 159 0.03 1.21 0.09448 0.00336 0.06176 0.00336 281.1 9.8 1 rim oz

25.1 358 0.07 9.14 0.03711 0.00298 0.12407 0.01812 174.7 14.0 1 rim oz

26.1 393 0.02 0.84 0.10419 0.00171 0.05872 0.00156 279.1 4.8 1 rim oz


E.1.22 SY9630 Syros Schist (Z2644, 97778)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 150 0.86 0.36 0.10444 0.00287 0.05424 0.00210 241.8 6.6 1 edge oz

1.2 227 1.19 2.66 0.11384 0.00507 0.07371 0.00407 270.7 11.9 1 edge oz

2.1 235 0.30 0.85 0.04212 0.00160 0.05616 0.00341 150.2 6.1 1 term oz

3.1 217 0.57 0.42 0.07973 0.00111 0.05304 0.00122 168.7 2.3 1 rim oz

4.1 113 0.81 0.32 0.10356 0.00203 0.05336 0.00138 219.7 4.3 1 centre oz

5.1 93 0.41 0.53 0.09461 0.00167 0.05448 0.00172 191.9 3.3 1 edge iz

6.1 135 0.78 0.48 0.11201 0.00116 0.05476 0.00160 221.8 2.3 1 rim oz

7.1 182 0.66 0.77 0.05100 0.00172 0.05582 0.00168 163.3 5.8 1 rim oz

7.2 160 0.74 0.38 0.10652 0.00136 0.05392 0.00150 221.7 2.8 1 core oxc

8.1 357 0.39 1.57 0.05350 0.00166 0.06208 0.00253 147.9 4.7 1 rim oz

9.1 468 0.24 1.18 0.04381 0.00103 0.05800 0.00210 106.9 2.6 1 edge oz

10.1 204 0.98 0.63 0.08822 0.00244 0.05607 0.00226 224.8 6.3 1 edge oz

11.1 156 0.29 1.76 0.03336 0.00303 0.06356 0.00332 145.9 13.4 1 edge oz

12.1 206 0.86 0.37 0.12243 0.00168 0.05354 0.00130 207.7 2.9 1 rim oz

13.1 540 0.54 1.08 0.04208 0.00262 0.05732 0.00155 113.5 7.1 1 term oz

14.1 204 0.33 0.92 0.06766 0.00091 0.05637 0.00111 132.1 1.8 1 rim oz

15.1 147 0.86 0.38 0.12753 0.00127 0.05388 0.00179 221.8 2.3 1 centre oz

16.1 69 1.84 2.64 0.07192 0.00150 0.07214 0.00225 208.7 4.8 1 centre iz

17.1 168 0.61 0.78 0.06388 0.00241 0.05682 0.00146 205.7 8.1 1 edge oz

18.1 181 0.64 0.32 0.10825 0.00158 0.05298 0.00175 203.6 2.9 1 centre oz

19.1 383 0.69 0.47 0.10818 0.00088 0.05417 0.00077 202.1 1.6 1 edge oz

20.1 292 0.56 0.48 0.08047 0.00096 0.05399 0.00094 187.2 2.4 1 edge oz

21.1 212 0.53 0.57 0.11650 0.00103 0.05533 0.00153 215.8 1.9 1 core oxc

E.1.23 NX9461 Naxos Calc-silicate (Z2298, 97779)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 819 0.15 29.30 0.03238 0.00512 0.28572 0.05684 111.9 17.6 1 term mix

2.1 645 0.15 4.15 0.05158 0.00151 0.08204 0.00655 120.6 3.5 1 term oz

3.1 103 0.23 1.94 0.23501 0.00292 0.07110 0.00341 431.5 6.4 1 edge iz

4.1 385 0.18 4.61 0.08526 0.00246 0.08669 0.00854 166.6 4.9 1 rim uol

Appendix E 291

5.1 87 0.18 3.62 0.25679 0.00668 0.08951 0.00541 622.8 15.5 1 core oz


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 426 1.44 3.62 0.03097 0.00052 0.07711 0.00418 74.6 1.3 1 edge oz

2.1 417 0.14 0.32 0.13520 0.00174 0.05515 0.00157 310.1 4.0 1 rim oz

2.2 269 0.61 1.32 0.11732 0.00193 0.06310 0.00246 304.1 4.9 1 core oxc

3.1 462 0.46 1.39 0.02502 0.00042 0.05890 0.00203 77.7 1.3 1 term oz

3.2 592 0.99 1.30 0.03102 0.00106 0.05812 0.00208 73.9 2.5 1 centre oz

4.1 22 0.28 3.00 0.38082 0.01097 0.09582 0.00572 991.4 26.6 1 core ixc

5.1 248 0.92 0.74 0.23604 0.00307 0.06506 0.00126 575.2 7.2 1 core oz

5.2 85 0.60 1.43 0.27378 0.00540 0.07066 0.00276 574.3 11.1 1 rim oz

6.1 118 0.76 2.81 1.56404 0.03576 0.20037 0.00254 2812.1 21.6 3 edge iz

6.2 113 0.32 8.63 1.29432 0.01616 0.24995 0.00454 3169.0 29.4 3 edge iz

7.1 2141 0.06 0.03 0.30502 0.00587 0.06036 0.00094 629.9 12.0 1 co edge iz

8.1 163 0.91 4.50 0.05875 0.00144 0.08605 0.00344 162.0 3.9 1 edge oz

9.1 332 0.27 0.66 0.26271 0.00605 0.06244 0.00122 499.3 11.6 1 term oz

10.1 752 0.08 0.18 0.14697 0.00140 0.05683 0.00092 431.2 4.2 1 term oz

11.1 721 0.25 1.18 0.02942 0.00060 0.05719 0.00218 74.5 1.5 1 rim oz

12.1 33 1.34 0.47 0.84619 0.01876 0.12783 0.00500 1842.6 106.5 3 core ixc

13.1 259 0.24 0.70 0.15038 0.00209 0.06148 0.00205 446.6 6.2 1 core oxc

14.1 121 0.72 3.36 0.05580 0.00132 0.07658 0.00495 153.9 3.6 1 term oz


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 622 0.32 2.86 0.04962 0.00136 0.07137 0.00420 96.7 2.6 1 rim uol

2.1 579 0.17 12.72 0.04766 0.00191 0.15160 0.01444 80.2 3.2 1 core oxc

3.1 539 0.34 4.11 0.07007 0.00160 0.08233 0.00303 136.0 3.1 1 term oo

4.1 97 0.19 4.81 0.39152 0.00475 0.10140 0.00391 685.3 7.9 1 core uxc

5.1 568 0.16 4.05 0.03939 0.00076 0.08076 0.00983 81.8 1.6 1 core oxc

5.2 1038 0.03 3.33 0.02800 0.00053 0.07433 0.00944 54.5 1.0 1 core oxc

6.1 324 0.19 4.71 0.07668 0.00151 0.08682 0.01519 114.7 2.3 1 edge oz


E.1.26 NX9464 Naxos Calc-Silicate (Z2038, 97782)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 565 1.08 0.13 0.25363 0.00207 0.06048 0.00056 577.0 4.5 1 in rim oz

2.1 619 0.24 1.60 0.11869 0.00235 0.06523 0.00112 290.5 5.6 1 rim oz

3.1 216 0.93 0.25 0.38934 0.00557 0.07033 0.00148 871.1 11.7 1 in rim oz

4.1 413 0.09 0.73 0.22396 0.00606 0.06471 0.00127 556.3 14.4 1 core oz

5.1 171 0.54 0.47 0.27503 0.00224 0.06519 0.00137 646.7 5.0 1 edge oz

6.1 956 0.12 0.30 0.09879 0.00116 0.05317 0.00079 225.2 2.6 1 edge oz

7.1 182 0.51 0.44 0.41222 0.00563 0.07508 0.00197 965.9 12.2 1 core oxc

8.1 132 0.51 0.61 0.42039 0.00548 0.07706 0.00142 982.8 11.9 1 core oxc

9.1 99 0.22 3.90 1.32874 0.01695 0.20788 0.00319 2859.2 28.0 3 core oxc

10.1 146 0.37 0.84 0.18730 0.00376 0.06296 0.00312 452.5 8.8 1 core oxc

11.1 235 0.89 0.73 0.24355 0.00350 0.06498 0.00166 564.4 7.8 1 rim oz

12.1 198 0.24 0.88 0.24839 0.00294 0.06671 0.00157 583.8 6.6 1 rim oz

13.1 175 0.13 2.39 0.16980 0.00210 0.07429 0.00299 401.5 4.8 1 centre oz

14.1 448 0.11 1.53 0.07543 0.00095 0.06238 0.00253 185.0 2.3 1 edge oz

15.1 36 1.03 2.22 0.77203 0.01244 0.12466 0.00365 1711.1 117.5 3 core oxc

16.1 268 0.33 1.64 0.74418 0.01291 0.11399 0.00138 1836.1 25.6 3 centre oz

17.1 293 0.51 0.28 0.76248 0.00941 0.10468 0.00114 1677.1 21.9 3 core oxc

18.1 317 0.65 1.34 0.12692 0.00169 0.06323 0.00261 296.8 3.9 1 rim oz

19.1 656 0.33 0.38 0.22771 0.00451 0.06151 0.00125 542.4 10.3 1 core oxc

20.1 79 0.48 3.11 0.24231 0.00703 0.08433 0.00611 566.8 15.7 1 centre oz

21.1 227 0.13 3.29 0.07556 0.00280 0.07661 0.00550 178.8 6.5 1 rim oz

1.1b 97 0.64 4.55 0.44766 0.01032 0.10546 0.00470 862.4 18.6 1 core oxc

1.2b 94 0.51 1.67 0.49620 0.00892 0.08505 0.00160 953.8 15.9 1 core oxc

2.1b 343 0.61 0.74 0.26115 0.00627 0.06668 0.00141 614.7 14.3 1 core oxc

2.2b 207 0.46 1.60 0.28534 0.00677 0.07430 0.00217 632.7 14.4 1 core oxc

3.1b 466 0.32 0.71 0.27515 0.00727 0.06669 0.00094 621.3 15.7 1 core oxc

3.2b 190 0.17 0.62 0.27874 0.00345 0.06687 0.00251 654.6 8.1 1 core oxc

4.1b 148 0.60 0.77 0.25876 0.00479 0.06608 0.00196 584.2 10.5 1 core oxc

4.2b 144 0.50 0.67 0.26960 0.00310 0.06585 0.00167 605.0 6.8 1 core oxc

5.1b 127 0.08 1.52 0.29726 0.00489 0.07404 0.00245 645.7 10.2 1 core oxc

5.2b 115 0.08 2.46 0.28086 0.00934 0.08113 0.00315 621.0 19.7 1 core oxc

6.1b 489 0.33 0.07 0.44138 0.00288 0.07079 0.00085 918.6 5.7 1 core oxc

Appendix E 293

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


7.1b 442 0.76 1.87 0.27167 0.00771 0.07434 0.00144 552.8 15.0 1 core oxc

8.1b 431 0.13 6.81 0.19576 0.00756 0.11064 0.00332 363.9 13.7 1 core oxc

9.1b 377 0.12 1.28 0.23928 0.00381 0.06666 0.00255 446.8 6.9 1 centre oz

10.1b 369 0.23 0.44 0.23445 0.00378 0.06049 0.00149 478.2 7.4 1 centre oz

10.2b 403 0.31 0.68 0.21749 0.00227 0.06159 0.00224 441.1 4.4 1 centre oz


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 186 0.01 17.38 0.01721 0.00144 0.18809 0.04791 33.0 2.7 1 rim uol

1.2 184 0.50 2.55 0.07023 0.00194 0.07160 0.00477 227.3 6.4 1 core oxc

2.1 1055 0.02 1.60 0.03080 0.00142 0.06019 0.00340 57.1 2.7 1 term oo

2.2 661 0.10 1.63 0.04811 0.00139 0.06270 0.00396 162.7 4.9 1 core uxc

3.1 232 0.55 1.20 0.09921 0.00319 0.06105 0.00415 243.1 7.7 1 core oxc

4.2 883 0.88 0.38 0.08156 0.00150 0.05394 0.00212 224.7 4.1 1 core oxc

5.1 870 0.36 0.73 0.07952 0.00086 0.05628 0.00263 201.2 2.1 1 edge oz

7.1 334 1.35 14.53 0.10444 0.00400 0.16902 0.00920 251.0 9.5 1 core oxc

8.1 143 0.23 5.22 0.03690 0.00135 0.09039 0.00995 94.8 3.4 1 centre iz

10.1 92 0.02 26.78 0.01392 0.00160 0.26453 0.05423 29.1 3.3 1 rim uol

10.2 218 0.52 2.45 0.09609 0.00242 0.07100 0.00857 236.6 5.9 1 core oxc

12.1 1106 0.01 6.75 0.02317 0.00093 0.10190 0.00987 48.0 1.9 1 rim mz


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


1.1 153 0.01 4.05 0.01340 0.00049 0.07966 0.00550 33.5 1.2 1 rim uol

1.2 232 0.02 8.36 0.00757 0.00036 0.11438 0.01308 15.0 0.7 1 rim uol

1.3 174 0.02 5.08 0.00615 0.00018 0.08766 0.00675 13.3 0.4 1 rim uol

1.4 144 0.06 3.28 0.05272 0.00098 0.07537 0.00460 136.3 2.5 1 rim uos

2.1 312 0.00 4.95 0.01868 0.00064 0.08711 0.00343 43.1 1.5 1 rim uos

4.1 293 0.13 1.22 0.09396 0.00382 0.06012 0.00103 203.5 8.2 1 term uos

5.1 290 0.01 3.74 0.02961 0.00065 0.07771 0.00364 65.7 1.4 1 edge rz

6.1 335 0.05 1.41 0.03587 0.00077 0.05889 0.00243 70.5 1.6 1 rim uol (?)


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


6.2 105 0.02 4.77 0.01391 0.00067 0.08544 0.01589 32.1 1.6 1 rim uol

7.1 172 0.32 1.42 0.12692 0.00404 0.06420 0.00227 314.8 9.8 1 rim oz

8.1 125 0.12 1.82 0.13080 0.00364 0.06800 0.00197 337.7 9.2 1 in rim iz

8.2 190 0.02 17.25 0.00896 0.00041 0.18668 0.02098 17.3 0.8 1 rim uol

8.3 133 0.06 10.25 0.05797 0.00433 0.13160 0.01073 120.2 8.9 1 rim uol

8.4 140 0.06 9.89 0.03999 0.00477 0.12854 0.01449 112.4 13.3 1 rim uol

9.1 535 0.01 3.68 0.02447 0.00076 0.07693 0.00519 50.2 1.6 1 rim uol (?)

10.1 212 0.24 1.79 0.10733 0.00104 0.06557 0.00224 239.8 2.3 1 edge oz

11.1 198 0.00 10.01 0.01767 0.00091 0.12817 0.01353 36.7 1.9 1 rim uol

11.2 421 0.01 5.82 0.02165 0.00066 0.09429 0.01254 46.2 1.4 1 rim uos

12.1 198 0.24 4.29 0.05170 0.00245 0.08425 0.00500 167.7 8.0 1 rim oo

13.1 258 0.01 17.44 0.03503 0.00172 0.18902 0.01556 62.5 3.1 1 rim uol

13.2 454 0.01 7.27 0.02974 0.00174 0.10637 0.01422 62.8 3.7 1 rim uol

14.1 127 0.04 5.95 0.03532 0.00079 0.09595 0.01469 78.1 1.7 1 rim oo

15.1 226 0.01 8.85 0.01531 0.00072 0.11872 0.01427 38.0 1.8 1 rim uol (?)

15.2 565 0.01 3.17 0.02222 0.00052 0.07280 0.00657 48.5 1.1 1 rim uol

16.1 201 0.01 7.87 0.02310 0.00067 0.11098 0.01506 48.1 1.4 1 rim uol

17.1 218 0.03 5.22 0.04157 0.00098 0.09037 0.01235 97.9 2.3 1 rim oo

18.1 327 0.00 4.50 0.01798 0.00107 0.08354 0.00910 45.1 2.7 1 rim uol

19.1 330 0.08 2.73 0.05970 0.00160 0.07070 0.00420 122.3 3.3 1 term uol

20.1 479 0.02 2.52 0.05108 0.00191 0.06859 0.00418 103.8 3.9 1 rim oo

21.1 455 0.04 1.43 0.04136 0.00047 0.05950 0.00195 93.0 1.1 1 rim oo

22.1 149 0.01 19.95 0.01542 0.00098 0.20899 0.04394 35.4 2.2 1 rim uol

22.2 177 0.07 7.46 0.02710 0.00216 0.10790 0.01347 62.3 4.9 1 rim uol

23.1 317 0.02 3.69 0.01821 0.00071 0.07696 0.00401 47.7 1.9 1 rim uol

24.1 312 0.04 3.13 0.05911 0.00329 0.07464 0.00642 156.4 8.6 1 term oz

25.1 203 0.55 0.67 0.20217 0.00361 0.06094 0.00208 427.3 7.5 1 edge iz

26.1 207 0.02 14.22 0.00707 0.00037 0.16198 0.04425 14.6 0.8 1 rim uol

27.1 196 0.02 22.86 0.00637 0.00055 0.23231 0.03220 13.5 1.2 1 rim uol

27.2 312 0.02 2.91 0.02878 0.00060 0.07092 0.00634 62.5 1.3 1 rim uol

28.1 218 0.02 3.61 0.06237 0.00259 0.07830 0.00514 148.0 6.1 1 term oz

29.1 370 0.01 11.74 0.02130 0.00061 0.14228 0.01734 41.4 1.2 1 core rz (?)

1.1a 139 0.02 17.82 0.00940 0.00054 0.19215 0.02347 29.2 1.7 1 rim uol

2.1a 79 0.03 15.19 0.01828 0.00274 0.17103 0.03553 52.7 7.9 1 term uol

3.1a 450 0.01 0.15 0.01724 0.00053 0.04815 0.00658 44.0 1.3 1 rim mzo

4.1a 169 0.01 8.09 0.01561 0.00068 0.11277 0.01199 35.1 1.5 1 rim uol

Appendix E 295

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


5.1a 143 0.07 6.25 0.04929 0.00189 0.09930 0.01505 115.5 4.4 1 rim oo

6.1a 138 0.25 2.91 0.08288 0.00247 0.07412 0.00500 212.7 6.2 1 term oz

7.1a 282 0.03 0.85 0.08143 0.00327 0.05649 0.00307 171.8 6.9 1 term oz

8.1a 140 0.01 7.98 0.04376 0.00181 0.11310 0.01132 101.1 4.2 1 rim oo

9.1a 315 0.09 0.57 0.06629 0.00122 0.05368 0.00478 146.9 2.7 1 rim oz

10.1a 156 0.04 10.73 0.01308 0.00116 0.13444 0.01629 40.5 3.6 1 edge rz

11.1a 1308 0.03 1.97 0.02861 0.00027 0.06340 0.00294 63.7 0.6 1 rim oz

12.1a 244 0.01 2.80 0.03414 0.00131 0.07051 0.00520 82.2 3.1 1 rim uo

13.1a 453 0.03 2.09 0.02841 0.00109 0.06458 0.00287 73.2 2.8 1 rim uol

13.2a 213 0.00 5.36 0.02080 0.00082 0.09077 0.01490 47.9 1.9 1 rim uol

16.1a 450 0.01 1.27 0.03215 0.00122 0.05780 0.00544 69.7 2.6 1 term uol

16.2a 153 0.01 25.76 0.01098 0.00117 0.25693 0.06442 28.1 3.0 1 term uol (?)

17.1a 535 0.01 0.17 0.02770 0.00080 0.04867 0.00298 62.8 1.8 1 term uol

17.2a 482 0.00 3.78 0.01758 0.00042 0.07783 0.00618 46.0 1.1 1 term oo

18.2a 169 0.01 10.45 0.00851 0.00092 0.13181 0.01831 22.0 2.4 1 rim uol

19.1a 207 0.00 4.78 0.01763 0.00051 0.08592 0.01198 40.9 1.2 1 term rz

19.2a 558 0.00 1.34 0.01839 0.00058 0.05795 0.00464 47.1 1.5 1 rim rz

20.1a 653 0.02 1.72 0.03739 0.00110 0.06200 0.00236 95.1 2.8 1 edge oz

22.1a 353 0.15 1.32 0.17629 0.00309 0.06384 0.00305 319.1 5.7 1 rim oo

22.2a 474 0.19 0.87 0.23206 0.00287 0.06284 0.00114 422.3 5.5 1 rim oo

23.1a 160 0.02 5.11 0.01579 0.00140 0.08817 0.01255 30.5 2.7 1 rim uol

23.2a 126 0.27 3.92 0.15648 0.00628 0.08463 0.00975 313.7 12.3 1 core oxc

24.1a 292 0.01 23.10 0.01484 0.00183 0.23420 0.03494 28.4 3.5 1 rim uol

24.2a 248 0.01 5.19 0.01342 0.00258 0.08883 0.01362 30.4 5.8 1 rim uol

25.1a 284 0.01 4.74 0.02066 0.00121 0.08527 0.00887 36.0 2.1 1 rim uol

26.1a 402 0.01 5.13 0.02617 0.00057 0.08871 0.00439 48.1 1.1 1 rim uol

27.1a 384 0.03 2.31 0.03292 0.00165 0.06615 0.00451 66.9 3.3 1 rim uol

28.1a 519 0.04 1.57 0.04884 0.00199 0.06067 0.00380 90.6 3.7 1 rim uol

29.1a 282 0.05 1.92 0.06827 0.00199 0.06476 0.00369 148.4 4.3 1 term oz

30.1a 165 0.01 28.23 0.00855 0.00100 0.27569 0.03605 14.1 1.7 1 rim oo

30.2a 266 0.01 31.15 0.02719 0.00147 0.29972 0.03859 38.9 2.1 1 rim uol

31.1a 179 0.08 2.34 0.06861 0.00219 0.06801 0.00514 142.4 4.5 1 rim oz

30.3a 371 0.49 0.00 0.39787 0.00483 0.06600 0.00139 764.0 9.1 1 core oxc

32.1a 158 0.02 6.17 0.01691 0.00114 0.09684 0.01696 35.1 2.4 1 term rz

33.1a 216 0.01 19.13 0.00706 0.00058 0.20175 0.02468 15.4 1.3 1 rim uol

34.1a 471 0.14 1.48 0.06404 0.00339 0.06082 0.00276 131.2 6.9 1 edge oz

35.1a 250 0.20 3.54 0.10728 0.00255 0.07879 0.00407 193.4 4.6 1 rim oz


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


36.1a 105 0.01 11.26 0.01330 0.00154 0.13797 0.04139 24.1 2.8 1 rim uol

36.2a 146 0.01 22.75 0.00788 0.00086 0.23117 0.04825 14.2 1.6 1 rim uol

36.3a 331 0.01 1.89 0.02239 0.00056 0.06237 0.00610 47.4 1.2 1 rim uol

37.1a 130 0.01 12.27 0.00818 0.00104 0.14609 0.03347 19.4 2.5 1 rim uol

37.2a 862 0.00 1.46 0.02282 0.00076 0.05872 0.00281 41.9 1.4 1 rim uol

38.1a 187 0.01 12.30 0.02202 0.00137 0.14696 0.02260 51.4 3.2 1 term rz

38.2a 209 0.03 0.98 0.03496 0.00158 0.05565 0.00318 81.4 3.7 1 term oz

39.2a 99 0.33 0.96 0.16812 0.00510 0.04846 0.00609 435.8 12.9 1 core oz

40.1a 172 0.02 1.05 0.00927 0.00082 0.05497 0.02224 20.1 1.8 1 rim uol

40.2a 213 0.02 5.78 0.00916 0.00055 0.09351 0.02145 23.7 1.4 1 rim uol

41.1a 231 0.02 2.51 0.02693 0.00059 0.02681 0.00600 55.2 1.2 1 term uol (?)

41.2a 177 0.04 3.39 0.03538 0.00199 0.02009 0.00346 74.9 4.2 1 rim uol

42.1a 138 0.63 1.82 0.71039 0.02399 0.12068 0.00273 1664.3 63.1 2 centre iz

43.1a 305 0.01 2.14 0.01877 0.00149 0.06421 0.00600 39.0 3.1 1 rim uol

44.1a 200 0.01 9.21 0.01103 0.00095 0.12132 0.01948 21.7 1.9 1 rim uol

1.1b 346 0.08 3.07 0.07326 0.00174 0.07457 0.00557 171.7 4.1 1 rim uol

1.2b 102 0.39 3.19 0.31568 0.01175 0.08654 0.01358 627.9 22.7 1 core oz

2.1b 852 0.00 3.63 0.02143 0.00047 0.07663 0.00971 49.1 1.1 1 rim uol

3.1b 301 0.05 3.45 0.05961 0.00099 0.07760 0.00660 169.9 2.8 1 edge uol

4.2b 506 0.01 3.06 0.02527 0.00064 0.07205 0.00786 53.5 1.4 1 rim oo

5.1b 1098 0.01 4.57 0.01590 0.00087 0.08419 0.00606 43.4 2.4 1 rim uol

5.2b 1478 0.01 2.39 0.02463 0.00058 0.06670 0.00509 56.3 1.3 1 rim uol

6.1b 380 0.72 0.64 0.18490 0.00252 0.06316 0.00105 522.3 6.8 1 in rim oo

7.1b 439 0.02 4.24 0.08416 0.00377 0.08438 0.00293 187.0 8.3 1 rim oo

8.1b 467 0.21 2.71 0.04265 0.00298 0.07112 0.00493 149.2 10.3 1 rim uol

8.2b 268 0.36 3.31 0.08921 0.00160 0.07828 0.00711 255.8 4.5 1 term oz

9.1b 236 0.03 19.16 0.01195 0.00068 0.20323 0.01863 39.4 2.3 1 rim uol

9.2b 95 0.20 5.28 0.04528 0.00296 0.09217 0.00891 155.9 10.1 1 rim uol

9.3b 214 0.31 1.00 0.10080 0.00170 0.06016 0.00215 282.9 4.7 1 centre oz

11.1b 599 0.03 4.54 0.02984 0.00067 0.08448 0.00758 69.7 1.6 1 edge uol

12.1b 556 0.03 5.78 0.01501 0.00091 0.09414 0.01473 44.2 2.7 1 rim uol

12.2b 68 3.20 1.87 0.22150 0.00605 0.07418 0.00550 563.3 14.8 1 core ixc

13.1b 159 0.02 26.80 0.00871 0.00105 0.26535 0.04827 21.3 2.6 1 rim mz

14.1b 230 0.01 14.19 0.01187 0.00070 0.16257 0.04087 33.0 2.0 1 rim mz

16.1b 167 0.01 20.91 0.01298 0.00145 0.21740 0.02584 32.7 3.6 1 rim mzo

16.2b 251 0.25 2.28 0.09369 0.00260 0.07129 0.00729 313.8 8.6 1 edge oxc

Appendix E 297

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


16.3b 682 0.09 2.12 0.04950 0.00109 0.06610 0.00141 138.6 3.0 1 rim mz

18.1b 112 0.37 5.61 0.08884 0.00419 0.09764 0.00450 284.6 13.2 1 edge oo

17.3b 173 0.12 0.04 0.05011 0.00115 0.04910 0.00499 163.6 3.7 1 edge oz

19.1b 179 0.01 13.05 0.01727 0.00505 0.15348 0.05908 47.4 13.8 1 term mz

20.1b 179 0.02 21.83 0.00978 0.00320 0.22497 0.07227 34.9 11.4 1 term uol

22.1b 177 0.01 10.75 0.01510 0.00085 0.13466 0.02469 41.7 2.3 1 rim uol

22.2b 412 0.01 6.34 0.01583 0.00090 0.09855 0.00910 37.3 2.1 1 rim uol

23.1b 613 0.02 18.52 0.03891 0.00172 0.19901 0.04252 97.3 4.3 1 rim oo

24.1b 360 0.00 12.49 0.01649 0.00113 0.14876 0.03374 37.1 2.5 1 term uol

25.1b 290 0.04 2.07 0.04168 0.00204 0.06550 0.00415 126.9 6.2 1 edge oo

26.1b 179 0.06 14.91 0.04416 0.00214 0.17000 0.01675 122.0 5.9 1 edge oz

27.1b 1006 0.04 12.55 0.09212 0.00193 0.15258 0.00429 218.1 4.5 1 edge oz

29.1b 441 0.00 13.29 0.02329 0.00106 0.15551 0.03501 48.8 2.2 1 rim mz

30.1b 133 0.07 22.02 0.04591 0.00229 0.22748 0.03024 95.6 4.7 1 rim uol

E.1.29 NX9490 Naxos Pelite (Z2264, 97781)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 73 1.32 0.40 1.77092 0.03681 0.25342 0.00398 3189.3 28.4 3 core ixc

1.2 56 1.12 0.60 1.24770 0.02900 0.17607 0.00487 2567.7 56.1 3 in rim iz

2.1 90 0.79 0.20 1.90096 0.04562 0.23466 0.01005 3072.1 68.0 3 core ixc

3.1 78 0.75 0.40 1.10891 0.03686 0.15277 0.00310 2336.2 46.3 3 core iz

4.1 740 0.74 1.05 0.33835 0.01001 0.07564 0.00269 813.1 22.6 1 term oz

4.2 169 0.71 0.42 0.46479 0.01163 0.07648 0.00287 1016.9 23.6 1 centre oz

5.1 101 0.93 0.50 1.08378 0.01206 0.13941 0.00160 2163.3 33.1 3 core iz

6.1 566 0.63 0.90 0.98852 0.02045 0.16866 0.00207 2469.6 33.6 3 rim sz

7.1 599 0.16 0.10 1.08569 0.02197 0.16177 0.00213 2464.1 22.6 3 term oo

7.2 59 0.65 0.50 1.30830 0.02820 0.16661 0.00243 2477.3 38.0 3 rim oo

8.1 68 0.67 1.92 0.38670 0.01165 0.08562 0.00297 874.2 24.6 1 core oxc

9.1 440 0.54 0.21 0.27973 0.00377 0.06320 0.00102 657.1 8.4 1 co edge oz

10.1 864 0.14 0.44 0.36998 0.00628 0.07019 0.00099 819.2 13.1 1 core oxc

10.2 246 0.59 0.46 0.40819 0.00868 0.07445 0.00233 946.6 18.7 1 core oxc

11.1 585 1.13 0.10 1.31220 0.02248 0.16858 0.00142 2553.9 14.5 3 core oxc

12.1 799 1.12 0.07 0.41150 0.00376 0.06705 0.00072 828.5 7.8 1 rim oz


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


12.2 637 1.26 0.42 0.31257 0.01217 0.06748 0.00112 738.1 27.2 1 core oz

13.1 1106 0.22 2.19 0.36146 0.01092 0.08503 0.00124 771.7 22.0 1 rim oz

13.2 204 0.65 0.40 0.81294 0.01013 0.10734 0.00174 1686.9 48.6 3 core oz

14.1 283 0.07 0.23 0.44965 0.00863 0.07368 0.00177 987.0 17.6 1 centre oz

15.1 636 0.09 1.44 0.10809 0.00295 0.07128 0.00290 527.8 16.5 1 rim oo

16.1 43 0.90 2.89 0.38199 0.00890 0.09753 0.00892 965.9 21.0 1 centre oz

17.1 120 1.01 1.91 0.25796 0.00686 0.07773 0.00253 608.1 15.4 1 centre oz

18.1 133 0.52 1.16 0.27481 0.00453 0.07132 0.00499 627.5 9.9 1 rim oz

19.1 416 0.81 0.65 0.42416 0.01207 0.07531 0.00186 920.0 24.5 1 core oz

19.2 1610 0.03 0.72 0.26554 0.00704 0.06640 0.00182 600.9 15.2 1 rim oo

20.1 281 0.34 0.20 1.14120 0.02427 0.12609 0.00222 2021.1 33.4 3 rim oz

21.1 58 0.78 2.09 0.34981 0.01213 0.08407 0.00482 771.1 25.2 1 centre oz

22.1 814 0.40 0.27 0.37255 0.00992 0.07043 0.00103 879.0 21.9 1 term oz

23.1 96 0.62 0.90 1.36113 0.04929 0.13285 0.00450 2023.4 86.3 3 core oz

24.1 220 0.91 0.00 1.21303 0.01970 0.16495 0.00258 2509.2 26.0 3 core oxc

25.1 179 0.56 0.30 1.22483 0.01694 0.13653 0.00162 2149.1 28.2 3 core ixc

26.1 293 0.61 0.20 0.88643 0.01624 0.12580 0.00127 2014.9 25.3 3 core oxc

27.1 188 0.54 1.04 0.37807 0.00576 0.07658 0.00224 845.7 12.1 1 core oz

28.1 164 1.07 0.60 0.90815 0.01097 0.11559 0.00240 1803.7 65.9 3 core ixc

1.1a 20 0.90 0.51 0.33353 0.00752 0.07253 0.00207 878.9 18.5 1 core oxc

2.1a 65 0.20 0.20 0.95571 0.01446 0.16259 0.00227 2467.0 24.1 3 core ixc

2.2a 149 0.13 0.28 0.22847 0.00802 0.06141 0.00173 571.7 19.3 1 rim oz

3.1a 45 0.66 0.40 0.43177 0.00498 0.07731 0.00107 1035.4 57.9 3 core ixc

3.2a 152 0.69 1.24 0.30984 0.00830 0.07693 0.00329 836.6 21.0 1 rim oz

4.1a 265 0.20 0.81 0.27772 0.00881 0.07132 0.00062 765.6 22.9 1 rim oz

1.1b 122 0.39 3.40 0.12643 0.00369 0.08177 0.00841 371.4 11.3 1 rim oz

2.1b 294 0.38 4.66 0.30233 0.00500 0.09801 0.00268 607.7 9.7 1 rim oz

3.1b 254 0.04 7.99 0.07076 0.00298 0.11350 0.02081 111.5 4.7 1 rim mix

4.1b 360 0.02 7.44 0.03962 0.00149 0.10875 0.01329 96.5 3.7 1 term mzo

5.1b 416 0.01 5.37 0.02446 0.00063 0.09088 0.00717 45.4 1.2 1 term mzo

6.1b 101 0.06 3.28 0.19532 0.00500 0.08122 0.00797 387.6 9.6 1 rim oo

7.1b 131 0.30 1.60 0.29906 0.00379 0.07672 0.00394 719.9 9.7 1 edge mix

8.1b 101 0.51 1.25 0.30627 0.00421 0.06962 0.00237 571.8 7.5 1 term oz

9.1b 219 0.01 0.43 0.33530 0.01011 0.06493 0.00317 638.4 18.3 1 term oz

Appendix E 299

E.1.30 NX94106 Naxos Pelite (Z2298, 97785)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 120 0.03 7.41 0.05482 0.00173 0.10896 0.00330 144.5 4.5 1 rim uol

1.2 142 0.10 13.27 0.05095 0.00054 0.15601 0.00447 121.6 1.3 1 rim uol

2.1 637 0.01 1.05 0.00579 0.00018 0.05492 0.00362 16.7 0.5 1 core oz

2.2 435 0.01 2.61 0.00647 0.00014 0.06758 0.00480 16.0 0.3 1 rim oz

3.1 335 0.64 0.27 0.24728 0.00260 0.06294 0.00055 629.3 6.3 1 core ixc

4.1 216 0.12 5.32 0.04761 0.00061 0.09181 0.00433 133.7 1.7 1 rim uol

5.1 177 0.00 1.85 0.00606 0.00017 0.06147 0.00534 18.8 0.5 1 rim uol

6.1 423 0.00 1.34 0.00649 0.00011 0.05730 0.00755 18.0 0.3 1 rim uol

7.1 147 0.01 3.01 0.00894 0.00034 0.07097 0.00784 24.2 0.9 1 rim uol

8.1 167 0.02 1.98 0.03429 0.00156 0.06396 0.00490 94.2 4.3 1 rim uol

9.1 210 0.02 2.60 0.01657 0.00041 0.06805 0.00410 45.8 1.1 1 rim uol

10.1 206 0.01 7.14 0.00631 0.00026 0.10440 0.01150 17.0 0.7 1 rim uol

11.1 203 0.01 2.96 0.01072 0.00053 0.07066 0.00520 31.2 1.5 1 rim mix

12.1 93 0.02 13.29 0.04233 0.00155 0.15609 0.00828 120.8 4.4 1 rim uol

13.1 304 0.02 2.46 0.01458 0.00044 0.06687 0.00270 42.5 1.3 1 rim uol

14.1 331 0.01 8.01 0.41879 0.00952 0.13714 0.00363 1044.3 21.9 1 rim iz

15.1 216 0.01 3.26 0.00617 0.00024 0.07284 0.01104 18.3 0.7 1 rim uos

15.2 175 0.00 5.26 0.00729 0.00040 0.08910 0.00858 18.1 1.0 1 rim uos

16.1 320 0.01 1.24 0.11966 0.00145 0.06217 0.00138 292.8 3.5 1 core oz

13.2 335 0.00 3.31 0.00710 0.00036 0.07329 0.01525 18.8 1.0 1 rim uol

17.1 222 0.01 10.17 0.00790 0.00024 0.12902 0.01106 20.1 0.6 1 rim uol

18.1 360 0.00 2.77 0.00696 0.00015 0.06888 0.00637 18.5 0.4 1 rim uol

20.1 337 0.01 3.33 0.00707 0.00033 0.07349 0.00525 20.0 0.9 1 rim uol

19.2 185 0.15 3.49 0.11358 0.00232 0.08068 0.00312 307.3 6.1 1 core oxc

21.1 143 0.01 1.14 0.04753 0.00159 0.05781 0.00243 127.2 4.2 1 rim uol

22.1 325 0.01 5.39 0.00662 0.00042 0.09018 0.00690 19.6 1.2 1 rim uol

23.1 370 0.02 6.04 0.03441 0.00195 0.09678 0.00689 89.4 5.0 1 rim mix

24.1 338 0.01 6.95 0.20322 0.00380 0.11407 0.00270 543.9 9.8 1 rim uol

25.1 153 0.01 9.34 0.00824 0.00055 0.12246 0.01097 29.5 2.0 1 rim uol

26.1 154 0.01 4.51 0.00711 0.00033 0.08309 0.00761 22.2 1.0 1 rim uol

27.1 126 0.01 3.29 0.00741 0.00036 0.07316 0.01119 21.9 1.0 1 rim uol


E.1.31 Ios Glaucophane Schist (Z2405, 89639)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


1.1 530 0.22 2.43 0.05976 0.00081 0.06883 0.00349 133.0 1.8 1 rim lm

1.2 331 0.24 2.18 0.10138 0.00174 0.06816 0.00369 198.1 3.4 1 term oo

2.1 472 0.15 1.46 0.13284 0.00178 0.06393 0.00271 277.6 3.6 1 rim oo

3.1 500 0.19 5.72 0.08186 0.00118 0.09696 0.00307 175.3 2.5 1 term oz

4.1 797 0.17 2.18 0.04683 0.00085 0.06627 0.00384 107.0 1.9 1 term mz

5.1 886 0.08 3.84 0.03128 0.00066 0.07926 0.00697 67.3 1.4 1 edge mz

6.1 499 0.11 8.44 0.07546 0.00618 0.11978 0.02033 188.7 15.2 1 edge oz

5.2 430 0.17 1.42 0.14292 0.00192 0.06340 0.00224 269.2 3.7 1 edge oz

7.1 116 0.17 12.12 0.06023 0.00289 0.15004 0.02678 178.4 8.7 1 edge oz

8.1 558 0.07 4.82 0.03137 0.00096 0.08766 0.00661 79.3 2.4 1 term oz

9.1 97 0.52 4.06 0.09731 0.00425 0.08648 0.00851 326.5 14.4 1 edge oz

10.1 692 0.08 6.56 0.02600 0.00137 0.10196 0.00989 76.2 4.0 1 edge mz

11.1 540 0.13 1.53 0.08709 0.00107 0.06260 0.00294 188.6 2.3 1 rim mzo

12.1 408 0.09 4.01 0.03832 0.00068 0.08096 0.00930 81.0 1.4 1 term mzo

13.1 329 0.11 2.69 0.07085 0.00148 0.07228 0.00612 192.0 4.1 1 term uol

14.1 350 0.12 1.46 0.12022 0.00179 0.06399 0.00236 278.0 4.1 1 edge mzo

15.1 366 0.17 1.52 0.10553 0.00175 0.06405 0.00209 258.7 4.3 1 edge oz

16.1 412 0.22 3.06 0.10711 0.00150 0.07573 0.00452 210.2 2.9 1 rim oz

17.1 350 0.19 4.50 0.07965 0.00346 0.08836 0.00908 244.0 10.6 1 edge oz

18.1 347 0.23 4.62 0.08560 0.00284 0.08783 0.00960 174.4 5.7 1 term mzo

19.1 495 0.12 9.96 0.05166 0.00176 0.13112 0.02636 125.7 4.3 1 edge oz

20.1 478 0.14 1.80 0.05662 0.00117 0.06363 0.00426 132.8 2.7 1 rim mzo

21.1 507 0.06 4.31 0.03431 0.00146 0.08357 0.01085 87.1 3.7 1 rim mzo

21.2 47 0.55 3.48 0.14245 0.00295 0.08137 0.00928 309.6 6.3 1 core oz

22.1 183 0.29 16.28 0.13419 0.00522 0.18559 0.02418 240.7 9.2 1 rim oz

22.2 263 0.26 1.12 0.12266 0.00208 0.06072 0.00226 259.6 4.3 1 core oz

23.1 429 0.30 2.95 0.10892 0.00220 0.07446 0.00670 193.7 3.9 1 edge rz

23.2 249 0.21 1.48 0.15946 0.00208 0.06532 0.00424 328.4 4.2 1 in rim oz

24.1 292 0.51 2.25 0.07404 0.00256 0.06900 0.00473 212.3 7.4 1 term oz

24.2 162 0.16 1.23 0.23295 0.00202 0.06719 0.00324 490.7 4.1 1 core oxc

25.1 586 0.09 5.01 0.03084 0.00058 0.08920 0.00966 78.7 1.5 1 rim mzo

26.1 619 0.09 16.77 0.03370 0.00148 0.18646 0.02550 60.0 2.6 1 rim mz

Appendix E 301

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


26.2 129 0.56 3.33 0.17051 0.00374 0.08010 0.00802 307.4 6.7 1 core oz

27.1 480 0.13 4.73 0.04260 0.00218 0.08736 0.00762 104.2 5.3 1 rim mz

27.2 67 0.78 6.39 0.15104 0.00290 0.10507 0.01728 297.8 5.6 1 core oz

28.1 492 0.11 4.85 0.04616 0.00104 0.08831 0.01121 102.5 2.3 1 rim mz

28.2 168 0.16 1.86 0.12635 0.00190 0.06720 0.00236 274.5 4.0 1 core oxc

29.1 110 0.27 4.02 0.12215 0.00513 0.08550 0.01258 295.1 12.1 1 term oz

29.2 118 0.42 2.71 0.14967 0.00185 0.07470 0.00748 295.5 3.6 1 core oxc

30.1 479 0.19 1.26 0.11214 0.00145 0.06140 0.00301 238.0 3.0 1 rim oz

31.1 453 0.13 2.41 0.06706 0.00189 0.06965 0.00529 177.5 5.0 1 term mzo

31.2 117 0.18 2.46 0.14778 0.00273 0.07401 0.00568 355.1 6.5 1 core oz

32.1 48 0.44 6.42 0.17037 0.00414 0.10572 0.02030 314.9 7.5 1 core oz

E.1.32 IO9615 Garnet-Glaucophane Schist (Z2644, 97786)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


1.1 235 0.05 2.30 0.02501 0.00237 0.06586 0.00266 60.5 5.7 1 rim uos

2.1 508 0.15 1.61 0.08186 0.00227 0.06166 0.00102 128.8 3.7 1 rim oz

3.1 331 0.06 0.23 0.37151 0.00901 0.06317 0.00053 647.1 15.2 1 rim oz

2.2 64 0.17 2.83 0.02469 0.00364 0.07106 0.00502 105.6 15.5 1 core oxc

4.1 204 0.22 1.88 0.05868 0.00278 0.06402 0.00382 136.2 6.4 1 rim uos

4.2 283 0.09 1.50 0.03340 0.00262 0.06029 0.00254 105.9 8.3 1 rim uos

5.1 91 0.01 3.87 0.00464 0.00025 0.07826 0.01267 43.2 2.8 1 rim uol

6.1 67 0.07 3.36 0.03977 0.00199 0.07681 0.00656 175.4 9.2 1 rim uos

7.1 321 0.03 5.16 0.02736 0.00081 0.08902 0.00305 57.7 1.7 1 rim oo

8.1 52 0.02 11.19 0.01082 0.00059 0.13761 0.01798 40.3 2.3 1 rim uol

9.1 176 0.02 3.47 0.03480 0.00067 0.07526 0.00588 54.7 1.1 1 rim uol

10.1 152 0.14 3.21 0.02988 0.00286 0.07368 0.00278 82.3 7.9 1 rim uol

11.1 38 0.07 3.61 0.02611 0.00284 0.07671 0.01053 71.8 7.8 1 term oz

12.1 157 0.33 3.29 0.04472 0.00146 0.07476 0.00571 102.4 3.3 1 rim oo

13.1 146 0.04 2.95 0.07682 0.00351 0.07313 0.00296 159.7 7.2 1 rim uol

14.1 401 0.03 1.70 0.02290 0.00084 0.06122 0.00299 72.6 2.7 1 term uol

15.1 91 0.07 23.19 0.00713 0.00079 0.23496 0.01531 41.4 4.8 1 rim uol

16.1 85 0.05 1.96 0.02895 0.00129 0.06350 0.00293 77.5 3.5 1 term uol

16.2 207 0.14 0.45 0.08827 0.00198 0.05393 0.00128 205.8 4.5 1 rim uol

17.1 175 0.01 11.17 0.00516 0.00090 0.13761 0.01802 48.4 8.7 1 rim uol


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


18.1 149 0.10 0.88 0.02390 0.00094 0.05471 0.00233 78.0 3.1 1 rim oo

19.1 158 0.03 6.72 0.01161 0.00111 0.10138 0.00470 44.3 4.3 1 rim uol

20.1 249 0.06 2.03 0.02028 0.00121 0.06359 0.00229 56.2 3.3 1 rim uol

21.1 116 0.01 5.98 0.00239 0.00019 0.09597 0.00932 74.8 7.7 1 rim uol

E.1.33 90346 Ios Quartz-Phengite Schist (Z2405, 90346)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


4.1 1135 0.04 10.59 0.01927 0.00094 0.13382 0.01200 61.7 3.0 1 rim uol

5.1 220 0.15 7.91 0.04290 0.00166 0.11455 0.00441 199.6 7.8 1 rim oz

6.1 589 0.02 2.35 0.04315 0.00053 0.06773 0.00202 127.2 1.6 1 rim mix

7.1 694 0.06 2.84 0.02426 0.00071 0.07081 0.00283 78.2 2.3 1 rim mzo

9.1 1155 0.04 2.17 0.01546 0.00064 0.06496 0.00248 61.8 2.5 1 term lm

10.1 693 0.09 1.69 0.02722 0.00033 0.06135 0.00164 77.2 1.0 1 rim oz

11.1 440 0.09 3.20 0.02416 0.00044 0.07372 0.00365 77.5 1.4 1 rim oo

12.1 574 0.04 3.65 0.03262 0.00092 0.07913 0.00234 165.7 4.9 1 term oz

13.1 227 0.46 4.33 0.23527 0.01209 0.10885 0.00113 1051.6 50.7 1 edge oz

14.1 237 0.21 1.72 0.04539 0.00137 0.06300 0.00199 144.3 4.3 1 term oz

15.1 620 0.11 3.02 0.03631 0.00161 0.07323 0.00347 128.3 5.6 1 term oz

16.1 340 0.11 8.47 0.03401 0.00127 0.11742 0.01357 114.0 4.2 1 term oz

18.1 438 0.12 5.43 0.09303 0.00186 0.09578 0.00147 241.0 4.7 1 term oz

19.1 337 0.18 4.30 0.05213 0.00082 0.08429 0.00516 137.9 2.2 1 term oz

19.2 294 0.20 1.58 0.08640 0.00115 0.06308 0.00220 197.2 2.8 1 core oz

20.1 454 0.07 3.96 0.05300 0.00193 0.08222 0.00426 174.5 6.3 1 rim oz

21.1 486 0.09 4.17 0.05042 0.00385 0.08348 0.00583 148.3 11.2 1 edge oz

22.1 312 0.06 1.88 0.09128 0.00242 0.06610 0.00108 221.0 5.8 1 centre oz

23.1 265 0.21 2.40 0.06831 0.00225 0.06998 0.00314 201.2 6.5 1 rim oz

24.1 280 0.17 4.32 0.06412 0.00116 0.08485 0.00379 155.4 2.8 1 edge oz

25.1 608 0.10 8.43 0.03514 0.00139 0.11810 0.00535 121.3 4.8 1 in rim oz

26.1 405 0.24 13.37 0.06547 0.00199 0.15958 0.01164 157.3 4.7 1 term oz

27.1 1222 0.16 5.07 0.04231 0.00127 0.09041 0.00383 126.8 3.8 1 term oz

28.1 868 0.01 7.67 0.02575 0.00042 0.11068 0.00755 60.4 1.0 1 rim uol

28.2 129 0.99 0.90 0.29734 0.00415 0.06981 0.00156 684.7 9.6 1 rim uol

29.1 630 0.06 5.61 0.04269 0.00096 0.09432 0.00878 96.9 2.2 1 term oz

Appendix E 303

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


30.1 193 0.08 7.40 0.06253 0.00146 0.11027 0.01223 157.5 3.7 1 term oz

31.1 427 0.04 7.72 0.05909 0.00181 0.11262 0.00274 144.0 4.4 1 term oz

32.1 483 0.02 2.71 0.05380 0.00106 0.07085 0.00310 121.2 2.5 1 term oz

33.1 59 1.84 1.12 0.39523 0.00711 0.08078 0.00242 972.4 16.4 1 centre oxc

35.1 578 0.05 5.10 0.03278 0.00123 0.08958 0.00387 71.0 2.7 1 term uol

36.1 439 0.02 5.09 0.02708 0.00035 0.08937 0.00658 61.9 0.9 1 rim uol

38.1 530 0.08 5.78 0.03129 0.00080 0.09568 0.00965 94.6 2.4 1 term uol

39.1 378 0.05 18.25 0.03748 0.00077 0.19822 0.02430 65.9 1.5 1 rim uol

40.1 182 0.18 14.95 0.05121 0.00404 0.17268 0.03392 162.5 12.7 1 rim uol

41.1 320 0.19 4.44 0.06113 0.00186 0.08642 0.00875 185.7 5.6 1 term oz

42.1 654 0.06 3.85 0.03114 0.00038 0.07943 0.00503 78.9 1.0 1 rim oz

43.1 362 0.08 11.85 0.04941 0.00314 0.14717 0.02476 167.5 10.6 1 term uol

44.1 639 0.19 8.88 0.08418 0.00870 0.12346 0.02590 208.7 21.2 1 rim oz

45.1 172 0.11 5.39 0.06479 0.00122 0.09597 0.00458 268.2 5.6 1 rim oz

E.1.34 FL9602 Folegandros Pelite (Z2633, 97787)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


1.1 397 0.03 0.20 1.09998 0.01621 0.12575 0.00193 2022.6 32.2 3 core oz

1.2 348 0.10 0.10 1.25809 0.02448 0.12718 0.00263 2044.1 38.0 3 rim oz

2.1 179 0.06 0.66 0.25516 0.00634 0.06152 0.00302 451.9 11.1 1 term oz

3.1 160 0.11 0.20 1.72922 0.02882 0.20821 0.00322 2880.1 26.2 3 rim oz

4.1 425 0.42 0.25 0.05279 0.00213 0.05305 0.00248 237.1 9.7 1 core oz

4.2 151 0.86 0.98 0.07838 0.00323 0.05943 0.00535 249.7 10.1 1 rim oz

4.3 92 1.95 2.54 0.06482 0.00351 0.07192 0.00689 215.6 11.5 1 rim oz

5.1 294 0.30 0.30 0.07346 0.00396 0.05794 0.00208 429.7 23.5 1 core oxc

5.2 213 0.57 0.75 0.09199 0.00220 0.05864 0.00176 299.2 7.0 1 rim oz

6.1 106 0.49 0.62 0.04189 0.00224 0.05442 0.00488 155.0 8.2 1 rim oz

6.2 97 0.78 1.36 0.05568 0.00248 0.06102 0.00382 172.1 7.6 1 core oz

7.1 743 0.26 0.57 0.06248 0.00168 0.05530 0.00198 218.0 5.8 1 rim oz

7.2 151 0.61 1.05 0.08685 0.00125 0.06019 0.00190 255.1 3.6 1 core oz

8.1 106 0.44 0.50 0.41273 0.01076 0.09191 0.00222 1373.3 84.1 3 core ixc

8.2 433 0.19 1.53 0.27761 0.01453 0.08185 0.00245 904.5 44.3 1 rim oz

9.1 607 0.37 5.53 0.01955 0.00275 0.09469 0.01628 91.3 12.8 1 term oz


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)


9.2 692 0.52 0.53 0.04438 0.00079 0.05330 0.00134 138.0 2.4 1 core iz

10.1 354 0.57 0.20 0.50487 0.01382 0.20801 0.00179 2879.9 15.2 3 core oz

11.1 736 0.61 0.75 0.06923 0.00111 0.05708 0.00177 229.9 3.7 1 rim uol

11.2 38 0.68 3.51 0.07802 0.00473 0.08038 0.00732 230.3 13.7 1 in rim oz

12.1 496 0.47 0.30 0.32396 0.04548 0.19082 0.00500 2732.6 47.6 3 term oz

13.1 644 0.19 0.58 0.14286 0.01110 0.05927 0.00115 744.8 55.6 1 rim oz

14.1 553 0.11 0.70 0.17194 0.00226 0.06311 0.00121 502.0 6.4 1 rim oz

15.1 226 0.29 8.89 0.22729 0.00535 0.13242 0.00114 529.6 12.1 1 rim oz

16.1 44 0.48 7.66 0.22729 0.00535 0.12231 0.00428 536.4 12.3 1 rim oz

17.1 123 0.16 1.41 0.22729 0.00535 0.07095 0.00255 571.2 13.0 1 rim oz

18.1 91 0.62 15.73 0.22729 0.00535 0.18875 0.00492 491.3 11.3 1 rim oz

19.1 550 0.10 1.43 0.22729 0.00535 0.07106 0.00104 571.1 13.0 1 term oz

E.1.35 SK9603 Sikinos Metabasic Schist (Z2633, 97788)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 194 0.52 1.07 0.10012 0.00118 0.05951 0.00198 224.3 2.6 1 core oz

1.2 305 0.62 0.22 0.08284 0.00112 0.05252 0.00219 225.6 3.2 1 in rim oz

1.3 270 0.54 3.91 0.07039 0.00170 0.08250 0.00856 200.9 5.0 1 edge oz

2.1 379 0.63 0.39 0.10544 0.00155 0.05382 0.00150 220.4 3.2 1 rim oz

2.2 516 0.47 0.73 0.07496 0.00181 0.05588 0.00186 186.1 4.5 1 rim oz

3.1 1196 0.42 3.49 0.03128 0.00089 0.07628 0.00343 65.3 1.9 1 edge oz

4.1 345 0.45 0.72 0.10312 0.00155 0.05650 0.00232 219.2 3.2 1 edge oz

4.2 788 0.45 1.62 0.08444 0.00112 0.06309 0.00583 180.4 2.4 1 edge oz

5.1 2033 0.48 22.11 0.03007 0.00052 0.23038 0.00905 46.5 0.8 1 edge iz

6.1 576 0.39 0.88 0.07574 0.00194 0.05707 0.00208 185.1 4.7 1 centre rz

7.1 365 0.35 0.99 0.10026 0.00133 0.05833 0.00130 201.1 2.7 1 rim oz

8.1 657 0.44 0.82 0.08117 0.00133 0.05659 0.00131 183.8 3.0 1 rim oz

9.1 1069 0.34 2.10 0.05778 0.00106 0.06616 0.00971 135.4 2.5 1 centre iz

10.1 691 0.63 0.86 0.10245 0.00219 0.05725 0.00180 200.8 4.3 1 rim oz

11.1 928 0.60 2.53 0.05665 0.00063 0.06959 0.00554 128.8 1.4 1 rim iz

1.1b 1029 0.47 0.85 0.03708 0.00188 0.05448 0.00090 69.5 3.5 1 centre oz

2.1b 717 0.37 1.41 0.07075 0.00074 0.06030 0.00174 128.4 1.3 1 edge oz

3.1b 575 0.70 0.32 0.07027 0.00259 0.05207 0.00088 162.6 6.0 1 rim oz

Appendix E 305

4.1b 344 0.41 1.25 0.12011 0.00222 0.06058 0.00196 201.5 3.7 1 centre oz

5.1b 327 0.58 0.36 0.14586 0.00164 0.05385 0.00107 227.6 2.8 1 edge oz

6.1b 429 0.51 1.02 0.10836 0.00232 0.05813 0.00089 177.1 3.8 1 rim oz

E.1.36 SIF9345 Sifnos Calc-silicate (Z2363, 97789)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.2 70 0.04 3.19 0.16500 0.00392 0.08133 0.01050 430.6 10.0 1 core oz

1.4 149 0.08 0.41 0.22909 0.00237 0.05939 0.00154 454.0 4.9 1 core oz

2.1 150 0.01 12.86 0.01923 0.00068 0.15150 0.04594 44.5 1.6 1 rim uos

2.2 108 0.41 4.48 0.22360 0.00794 0.09426 0.00661 534.7 18.2 1 core iz

3.1 981 0.00 2.94 0.01266 0.00045 0.07068 0.00485 37.9 1.4 1 rim lm

4.1 412 0.01 15.31 0.01987 0.00087 0.17148 0.02917 45.2 2.0 1 rim lm

5.1 444 0.06 7.23 0.05501 0.00284 0.10674 0.02079 97.5 5.0 1 rim uol

5.2 111 0.53 0.78 0.08742 0.00138 0.05689 0.00294 217.5 3.4 1 core oz

6.1 302 0.01 14.47 0.02157 0.00136 0.16468 0.02826 46.9 2.9 1 core ixc

7.1 329 0.02 8.15 0.04115 0.00174 0.11379 0.01619 76.5 3.2 1 rim uol

7.2 179 0.32 11.41 1.02393 0.01535 0.18964 0.00257 2667.7 28.2 2 core oxc

8.1 682 0.10 2.37 0.05321 0.00159 0.06796 0.00423 132.4 3.9 1 rim uos

9.1 935 0.01 4.98 0.01694 0.00050 0.08735 0.01647 42.0 1.2 1 rim uos

9.2 480 0.68 0.26 0.11471 0.00084 0.05413 0.00064 286.6 2.1 1 rim uos

10.1 338 0.02 14.16 0.01407 0.00101 0.16193 0.01414 34.5 2.5 1 rim uos

12.1 213 0.07 31.59 0.02034 0.00305 0.30375 0.07218 38.9 5.8 1 term uol

12.2 232 1.06 0.75 0.11612 0.00195 0.05840 0.00150 295.3 4.9 1 core iz

13.1 312 0.07 16.77 0.04439 0.00414 0.18434 0.04244 106.4 9.8 1 rim lm

14.1 379 0.03 8.50 0.06346 0.00196 0.11726 0.02068 110.4 3.5 1 rim uos

15.1 625 0.01 8.91 0.01489 0.00070 0.11926 0.02466 39.3 1.9 1 core iz

16.2 438 0.07 8.87 0.04079 0.00102 0.11998 0.01522 92.2 2.3 1 rim uol

17.1 147 0.38 3.93 0.09714 0.00293 0.08472 0.01374 322.3 9.9 1 edge oz

17.2 259 0.38 9.40 0.16208 0.00398 0.13020 0.01650 384.2 9.2 1 term oz

18.1 321 0.01 16.33 0.02412 0.00094 0.17986 0.02948 51.7 2.0 1 rim uos

20.2 642 0.01 7.04 0.03013 0.00126 0.10440 0.01669 56.0 2.3 1 rim uos

21.1 999 0.01 11.45 0.02217 0.00054 0.14017 0.02571 50.7 1.2 1 rim uos

21.3 819 0.31 0.12 0.21596 0.00134 0.05986 0.00085 561.1 3.7 1 core oz

22.1 376 0.01 11.43 0.03130 0.00092 0.14004 0.03129 54.9 1.6 1 rim uos

25.1 319 0.03 33.70 0.02230 0.00252 0.32104 0.04343 45.6 5.1 1 rim uos

26.1 263 0.04 16.25 0.02191 0.00170 0.17920 0.04736 53.5 4.1 1 rim uos


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


27.1 292 0.01 20.56 0.02802 0.00140 0.21418 0.05975 48.2 2.4 1 rim uos

28.1 63 0.64 1.18 0.09459 0.00130 0.06062 0.00277 238.8 3.2 1 core oz

29.1 384 0.18 0.24 0.09704 0.00094 0.05396 0.00098 286.2 3.1 1 core oz

30.1 200 0.12 3.48 0.26721 0.00292 0.08732 0.00095 568.5 6.0 1 core oz

31.1 187 0.33 0.38 0.26136 0.00201 0.06340 0.00136 613.5 4.5 1 rim oz

32.1 472 0.41 0.22 0.13516 0.00106 0.05439 0.00123 310.0 2.4 1 in rim oz

33.1 236 0.60 0.26 0.11687 0.00101 0.05470 0.00102 311.3 2.8 1 core uxc

33.2 464 0.41 0.91 0.12788 0.00138 0.05865 0.00283 247.9 2.9 1 rim oz

34.1 168 0.17 3.29 0.11071 0.00310 0.08042 0.00686 350.7 10.0 1 core uxc

35.1 524 0.13 2.57 0.04292 0.00061 0.06866 0.00754 73.8 1.2 1 rim uol

36.1 420 0.71 11.14 0.02301 0.00065 0.13900 0.01859 55.2 1.6 1 rim uol

37.1 128 0.42 3.98 0.60794 0.01023 0.11920 0.00255 1692.6 58.8 2 core uxc

38.1 460 0.23 1.71 0.08074 0.00092 0.06341 0.00216 163.6 1.9 1 core oxc

39.1 362 0.87 0.88 0.06204 0.00085 0.05622 0.00191 147.0 2.0 1 in rim oz

40.1 162 0.37 0.67 0.05265 0.00080 0.05435 0.00370 141.9 2.2 1 rim oz

41.1 90 0.16 0.49 0.20781 0.00241 0.06241 0.00179 545.9 6.3 1 core oxc

42.1 69 0.39 4.06 0.56651 0.01013 0.13291 0.00326 2021.5 62.9 2 core oxc

43.1 218 0.56 3.58 0.16003 0.00187 0.08505 0.00399 444.1 5.5 1 in rim oz

44.1 132 0.58 0.24 0.19455 0.00286 0.05811 0.00112 458.9 6.5 1 in rim oz

E.1.37 89642 Syros Retrogressed Eclogite (Z2405, 89642)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.2 56 0.58 4.01 0.03322 0.00283 0.08117 0.01730 93.2 7.9 1 centre oz

1.3 123 0.79 19.05 0.03159 0.00211 0.20545 0.01894 75.6 5.0 1 term oz

2.2 91 0.44 0.56 0.02723 0.00249 0.04330 0.01455 92.2 8.4 1 core oz

3.3 299 0.81 5.50 0.02926 0.00116 0.09316 0.01136 74.4 2.9 1 term oz

4.1 293 1.22 6.21 0.03124 0.00117 0.09913 0.01360 79.8 3.0 1 core oz

5.1 53 0.41 27.72 0.04052 0.00410 0.27756 0.02774 93.1 9.4 1 term oz

5.2 117 0.82 16.55 0.03178 0.00277 0.18476 0.02322 74.7 6.5 1 co edge oz

5.3 206 0.92 12.03 0.04288 0.00282 0.14785 0.03527 106.6 6.9 1 term oz

5.4 117 0.80 33.00 0.03405 0.00203 0.32098 0.02556 70.7 4.2 1 co edge oz

9.2 93 0.47 59.59 0.08003 0.00594 0.54156 0.04218 87.1 6.4 1 core uxc

9.3 27 0.46 33.79 0.04015 0.00156 0.32354 0.02675 76.8 3.0 1 rim oz

Appendix E 307

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


10.3 240 0.66 29.29 0.04563 0.00289 0.29049 0.02164 85.5 5.4 1 term oo

11.1 44 0.49 57.88 0.07070 0.00441 0.52737 0.04211 85.1 5.3 1 rim oz

12.1 285 0.95 17.61 0.03107 0.00142 0.19342 0.01619 68.8 3.1 1 term oz

12.2 270 0.87 15.56 0.02635 0.00171 0.17643 0.01313 71.2 4.6 1 term oz

12.3 32 0.72 46.11 0.04999 0.00512 0.42420 0.06706 76.8 7.8 1 term oz

13.1 56 0.41 59.67 0.05835 0.00503 0.54205 0.04411 66.8 5.7 1 edge oz

13.3 21 0.85 27.04 0.03550 0.00263 0.26839 0.03214 72.3 5.3 1 core oz

14.1 396 0.87 14.52 0.03231 0.00244 0.16787 0.01259 74.9 5.6 1 term lol

14.2 34 0.44 38.13 0.04659 0.00329 0.36380 0.04026 92.8 6.5 1 core uxc

14.3 10 0.63 48.85 0.05551 0.00716 0.44659 0.06910 79.0 10.1 1 core oz

15.1 111 0.77 17.25 0.03311 0.00153 0.19051 0.02397 75.2 3.5 1 rim oz

16.1 146 0.70 18.77 0.03315 0.00217 0.20312 0.01965 73.1 4.8 1 edge oz

E.1.38 NX9301 Naxos I-type Granodiorite (Z1870, 97790)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


4.1 228 0.31 30.88 0.00717 0.00030 0.29696 0.02641 12.5 0.6 1 term oz

5.1 238 0.25 19.85 0.00585 0.00015 0.20739 0.01099 12.1 0.4 1 rim oz

6.1 230 0.23 21.25 0.00585 0.00016 0.21881 0.01100 12.0 0.4 1 term oz

7.1 229 0.20 17.57 0.00561 0.00014 0.18890 0.00775 11.9 0.4 1 rim oz

8.1 216 0.20 21.09 0.00581 0.00017 0.21751 0.01055 12.4 0.4 1 rim oz

8.2 195 0.24 12.59 0.00524 0.00011 0.14852 0.00455 12.6 0.3 1 term oz

9.1 151 0.36 26.82 0.00641 0.00014 0.26401 0.01861 12.8 0.3 1 rim oz

10.1 273 0.24 23.84 0.00669 0.00025 0.23977 0.01204 11.9 0.6 1 edge oz

11.1 233 0.24 12.93 0.00510 0.00016 0.15126 0.01001 12.4 0.4 1 rim oz

14.1 310 0.44 13.76 0.00560 0.00021 0.15795 0.00660 13.0 0.5 1 rim oz

15.1 297 0.18 14.34 0.00464 0.00012 0.16272 0.01121 12.2 0.3 1 rim oz

16.1 182 0.20 25.46 0.00585 0.00021 0.25294 0.01467 11.6 0.5 1 term oz

17.1 270 0.21 17.79 0.00475 0.00011 0.19069 0.01100 12.1 0.3 1 rim oz

18.1 282 0.16 14.16 0.00475 0.00008 0.16124 0.01048 12.4 0.2 1 rim oz

19.1 198 0.61 21.08 0.00548 0.00012 0.21739 0.02379 13.0 0.3 1 term oz

20.1 230 0.21 19.74 0.00525 0.00022 0.20655 0.01230 12.4 0.5 1 term oz

21.1 205 0.27 22.44 0.00486 0.00010 0.22846 0.01142 13.0 0.3 1 rim oz

22.1 308 0.22 10.79 0.00548 0.00010 0.13386 0.00885 12.6 0.3 1 term oz

24.1 276 0.25 16.72 0.00503 0.00013 0.18204 0.00797 12.6 0.3 1 term oz


Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


25.1 188 0.53 26.01 0.00513 0.00008 0.25746 0.01538 12.7 0.2 1 term oz

26.1 311 0.25 16.15 0.00568 0.00014 0.17743 0.00737 12.8 0.4 1 rim oz

27.1 233 0.26 25.25 0.00570 0.00022 0.25127 0.01930 12.9 0.5 1 term oz

E.1.39 NX9303 Naxos Fractionated I-type Granite (Z2298, 97791)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 328 0.26 1.92 0.00407 0.00009 0.06185 0.00585 12.3 0.3 1 term oz

2.1 346 0.16 2.52 0.00365 0.00011 0.06667 0.00331 11.1 0.3 1 term oz

3.1 144 0.51 2.66 0.00432 0.00015 0.06787 0.00472 12.2 0.4 1 centre oz

3.2 247 0.27 3.26 0.00376 0.00011 0.07275 0.00416 12.0 0.4 1 term oz

4.1 195 0.24 13.77 0.00422 0.00019 0.15806 0.00906 10.8 0.5 1 edge oz

5.1 269 0.38 0.67 0.00483 0.00033 0.05178 0.00783 13.8 1.0 1 edge oz

6.1 173 0.32 3.00 0.00427 0.00013 0.07064 0.01220 12.2 0.4 1 rim oz

7.1 248 0.45 0.51 0.00400 0.00017 0.05042 0.00585 12.3 0.5 1 edge oz

8.1 268 0.24 0.14 0.00520 0.00116 0.04738 0.00360 10.2 2.3 1 rim oz

9.1 250 0.21 5.92 0.00386 0.00015 0.09429 0.01108 11.2 0.4 1 rim oz

10.1 99 0.50 9.57 0.00385 0.00013 0.12391 0.01381 10.8 0.4 1 rim oz

11.1 164 0.17 7.27 0.00394 0.00013 0.10528 0.01280 11.3 0.4 1 core oz

E.1.40 NX9470 Naxos I-type Granitoid (Z2613, 97792)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 43 0.25 0.33 0.11661 0.00247 0.05405 0.00092 261.0 5.4 1 core oxc

1.2 593 0.04 10.54 0.00878 0.00025 0.13202 0.01197 16.9 0.5 1 rim mzo

2.1 1261 0.11 12.24 0.01600 0.00037 0.14585 0.01102 18.8 0.5 1 term oz

3.1 216 0.37 0.64 0.00663 0.00021 0.05156 0.00155 14.7 0.5 1 edge oz

4.1 327 0.60 0.65 0.00556 0.00013 0.05159 0.00184 14.9 0.3 1 edge oz

6.1 106 0.56 14.53 0.00576 0.00020 0.16431 0.01235 13.3 0.5 1 edge oz

7.1 282 0.13 0.74 0.00856 0.00024 0.05234 0.00216 16.0 0.5 1 term oz

9.1 962 0.04 0.00 0.00805 0.00004 0.04634 0.00101 15.5 0.1 1 edge oz

10.1 84 0.31 3.31 0.00512 0.00032 0.07321 0.00665 15.1 0.9 1 edge oz

Appendix E 309

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


11.1 16 0.56 0.68 0.00563 0.00007 0.05182 0.00147 13.3 0.2 1 edge oz

12.1 246 0.65 0.46 0.00650 0.00008 0.05000 0.00141 13.3 0.2 1 core oz

13.1 211 0.72 0.29 0.00598 0.00013 0.04866 0.00232 13.4 0.3 1 core oz


E.1.41 NX9446 Naxos S-type Granite (Z2613, Z2644, 97793)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 189 0.14 0.66 0.00419 0.00020 0.05165 0.00344 12.5 0.6 1 centre oz

2.1 146 0.46 0.65 0.00462 0.00008 0.05161 0.00292 13.8 0.2 1 centre oz

3.1 171 0.02 0.88 0.01020 0.00013 0.05361 0.00168 23.4 0.3 1 in rim oz

4.1 139 0.24 0.38 0.08452 0.00083 0.05253 0.00194 169.9 1.7 1 rim oz

5.1 185 0.78 0.34 0.16756 0.00384 0.05521 0.00215 301.8 6.9 1 core sz

6.1 540 0.07 1.44 0.01037 0.00032 0.05805 0.00317 17.4 0.5 1 core oxc

7.1 779 0.07 1.09 0.00725 0.00016 0.05516 0.00357 13.6 0.3 1 core oz

8.1 370 0.32 2.57 0.00692 0.00030 0.06714 0.00287 13.8 0.6 1 term oz

9.1 173 0.03 8.25 0.00501 0.00221 0.11363 0.06498 34.0 15.0 1 rim uol

10.1 1272 0.30 0.84 0.00688 0.00012 0.05313 0.00198 12.9 0.2 1 centre oz

11.1 270 0.29 5.72 0.00633 0.00017 0.09267 0.00991 12.0 0.3 1 edge oz

12.1 894 0.34 0.98 0.00746 0.00011 0.05429 0.00230 13.0 0.2 1 rim oz

13.1 590 0.51 1.04 0.00538 0.00015 0.05471 0.00327 11.9 0.3 1 rim oz

14.1 292 0.34 3.05 0.00601 0.00013 0.07107 0.00588 12.4 0.3 1 rim oz

15.1 537 0.45 2.03 0.00751 0.00015 0.06277 0.00302 12.3 0.3 1 rim oz

16.1 503 0.05 2.33 0.00619 0.00009 0.06525 0.00340 14.1 0.2 1 term oz

17.1 325 0.14 0.12 0.12570 0.00299 0.05366 0.00112 313.0 7.3 1 term oz

18.1 895 0.23 1.35 0.00637 0.00007 0.05729 0.00283 13.4 0.2 1 term oz

19.1 474 0.27 1.55 0.00510 0.00007 0.05891 0.00297 13.0 0.2 1 rim oz

E.1.42 TIN9603 Tinos S-type Granite (Z2665, 97794)

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


1.1 667 0.60 1.27 0.00662 0.00024 0.05686 0.00309 14.3 0.5 1 rim oz

2.1 1142 0.39 1.15 0.00693 0.00026 0.05589 0.00456 14.3 0.5 1 term oz

3.1 3809 0.13 0.47 0.00685 0.00034 0.05025 0.00251 14.9 0.7 1 term oz

4.1 501 0.56 4.53 0.00654 0.00027 0.08402 0.00781 14.3 0.6 1 rim oz

5.1 366 0.31 4.91 0.00597 0.00032 0.08717 0.00918 13.3 0.7 1 term oz

6.1 836 0.45 2.61 0.00718 0.00027 0.06802 0.00642 14.7 0.5 1 term oz

7.1 831 0.22 4.07 0.00796 0.00041 0.08021 0.00519 15.8 0.8 1 rim oz

8.1 364 0.49 6.67 0.00634 0.00019 0.10177 0.01226 13.1 0.4 1 edge oz

Appendix E 311

Spot U

(ppm)

Th/U f % 206Pb /238U


±1σσσσ Age

(Ma)


9.1 534 0.40 5.03 0.00803 0.00032 0.08814 0.01272 14.8 0.6 1 term oz

10.1 378 0.17 6.18 0.00707 0.00029 0.09775 0.00698 14.1 0.6 1 edge oz

11.1 665 0.50 2.89 0.00706 0.00019 0.07034 0.00496 14.6 0.4 1 edge oz

12.1 244 0.30 10.33 0.00764 0.00027 0.13228 0.02083 14.8 0.5 1 centre oz

13.1 802 0.63 2.56 0.00682 0.00032 0.06765 0.00648 14.8 0.7 1 edge oz

14.1 792 0.41 3.67 0.00836 0.00023 0.07691 0.00966 15.3 0.4 1 rim oz

15.1 949 0.19 2.74 0.00773 0.00013 0.06911 0.00369 14.5 0.2 1 term oz

16.1 819 0.21 2.07 0.00697 0.00018 0.06353 0.00315 13.9 0.4 1 rim oz

17.1 521 0.17 5.72 0.00660 0.00022 0.09388 0.00868 12.9 0.4 1 edge oz


E.2 Monazite U-Th-Pb Analytical Results

All ages are calculated from 208Pb/232Th ratios.

E.2.1 NX9637 Naxos Melt Pod Naxos Migmatite (Z2922, 97774)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 208Pb /

232Th

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

1.1 28351 4 0.23 0.08741 0.00170 0.05195 0.00072 0.05195 0.00072 14.8 0.5

2.1 24258 5 0.48 0.14668 0.00247 0.04544 0.00154 0.05727 0.000972 13.1 0.5

3.1 15810 9 0.16 0.10406 0.00071 0.05727 0.00097 0.04979 0.00062 13.4 0.3

4.1 32120 5 0.16 0.11371 0.00054 0.04979 0.00062 0.04734 0.00071 14.2 0.3

5.1 25946 5 0.19 0.09955 0.00222 0.04734 0.00071 0.04544 0.00154 12.9 0.4

6.1 23220 5 0.53 0.12098 0.00330 0.05106 0.00201 0.05106 0.00201 12.8 0.5

8.1 27120 4 0.54 0.12497 0.00204 0.05293 0.00154 0.05293 0.00154 13.0 0.4

9.1 13677 6 0.23 0.11331 0.00316 0.05725 0.00105 0.05725 0.00105 13.0 0.5

10.1 26516 3 2.24 0.03564 0.00192 0.11061 0.00616 0.11061 0.00616 11.0 0.8

11.1 16753 9 0.15 0.08947 0.00064 0.05590 0.00101 0.05590 0.00101 12.6 0.3

12.1 25502 6 0.30 0.10410 0.00206 0.05089 0.00120 0.05090 0.00120 12.3 0.4

13.1 7934 22 0.12 0.09446 0.00097 0.07713 0.00174 0.07713 0.00174 10.8 0.3


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 208Pb /

232Th

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

1.1 11610 7 0.19 0.12763 0.00189 0.05622 0.00106 0.05622 0.00106 15.7 0.5

2.1 11970 8 0.18 0.12426 0.00159 0.05357 0.00103 0.05360 0.00103 15.3 0.4

3.1 7128 14 0.26 0.09909 0.00221 0.05945 0.00254 0.05945 0.00254 12.1 0.5

4.1 9574 11 0.32 0.09260 0.00515 0.06701 0.00299 0.06701 0.00299 12.7 0.9

5.1 15325 7 0.18 0.11255 0.00214 0.05522 0.00097 0.05522 0.00097 13.9 0.4

6.1 15838 8 0.16 0.11024 0.00174 0.05411 0.00094 0.05411 0.00094 14.1 0.4

7.1 9310 9 0.18 0.12162 0.00105 0.05746 0.00122 0.05746 0.00122 14.6 0.4

8.1 13599 7 0.18 0.12278 0.00088 0.05608 0.00100 0.05609 0.00100 14.8 0.4

9.1 13758 8 0.18 0.08851 0.00069 0.05544 0.00110 0.05544 0.00110 12.7 0.2

10.1 13920 8 0.17 0.10717 0.00138 0.05611 0.00105 0.05611 0.00105 12.9 0.4

11.1 11638 6 0.24 0.11458 0.00232 0.05745 0.00114 0.05745 0.00114 12.8 0.4

12.1 14963 7 0.20 0.09348 0.00117 0.05581 0.00105 0.05581 0.00105 12.5 0.3

Appendix E 313


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 208Pb /

232Th

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

1.1 14149 9 0.16 0.12045 0.00083 0.06923 0.00109 0.06923 0.00109 15.4 0.3

2.1 13130 9 0.15 0.13649 0.00097 0.05794 0.00100 0.05794 0.00100 14.9 0.5

3.1 16321 7 0.19 0.10248 0.00133 0.05782 0.00099 0.05782 0.00100 13.5 0.3

9.1 23126 6 0.18 0.09604 0.00110 0.05591 0.00083 0.05592 0.00083 13.4 0.3

10.1 15048 8 1.23 0.11800 0.00133 0.07952 0.00733 0.07952 0.00733 13.4 0.3

11.1 17132 8 0.17 0.09665 0.00066 0.05539 0.00095 0.05539 0.00095 13.0 0.3

12.1 28649 6 0.17 0.10262 0.00134 0.05643 0.00074 0.05643 0.00074 12.8 0.3

13.1 28310 6 0.17 0.11331 0.00058 0.05431 0.00070 0.05431 0.00070 13.5 0.3

14.1 11065 12 0.14 0.10864 0.00149 0.05808 0.00117 0.05808 0.00117 13.6 0.3

15.1 15979 8 0.16 0.11314 0.00196 0.05687 0.00096 0.05687 0.00096 12.9 0.4

16.1 16526 8 0.16 0.11195 0.00075 0.05497 0.00092 0.05497 0.00092 13.6 0.3


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 208Pb /

232Th

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

1.1 8538 11 0.16 0.12914 0.00303 0.05902 0.00125 0.05902 0.00125 15.9 0.6

2.1 9746 10 0.33 0.14625 0.00113 0.16437 0.00206 0.16437 0.00206 15.0 0.4

3.1 9488 10 0.16 0.12078 0.00101 0.06286 0.00125 0.06286 0.00125 16.2 0.4

4.1 8256 11 0.17 0.11967 0.00250 0.06406 0.00140 0.06406 0.00140 14.3 0.5

5.1 7766 11 0.17 0.11291 0.00201 0.06294 0.00146 0.06294 0.00146 14.0 0.4

6.1 4958 13 0.20 0.10672 0.00456 0.07555 0.00217 0.07555 0.00217 13.0 0.8

7.1 4633 18 0.22 0.10687 0.00136 0.10772 0.00264 0.10772 0.00264 12.3 0.3

7.1 8432 9 0.30 0.10921 0.00177 0.06530 0.00143 0.06530 0.00143 12.9 1.0

9.1 10259 18 0.47 0.10824 0.01094 0.06596 0.00523 0.06597 0.00523 12.1 1.6

10.1 8556 12 0.15 0.11250 0.00104 0.05966 0.00134 0.05966 0.00134 14.1 0.4

11.1 5633 13 0.41 0.11839 0.00136 0.06297 0.00389 0.06297 0.00389 12.9 0.4

12.1 8156 15 0.13 0.10268 0.00238 0.06583 0.00150 0.06583 0.00150 13.4 0.5


E.2.5 NX9438 Naxos Pegmatite (Z2301, 97798)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 208Pb /

232Th

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

1.1 10163 8 0.20 0.07648 0.00140 0.06149 0.00130 0.06149 0.00130 16.9 0.4

2.1 8674 13 0.15 0.07745 0.00191 0.06653 0.00142 0.06925 0.00160 17.0 0.5

3.1 7587 12 0.17 0.07977 0.00128 0.06925 0.00161 0.07598 0.00253 17.0 0.4

4.1 6475 13 0.27 0.08189 0.00132 0.07598 0.00253 0.06803 0.00217 16.8 0.4

5.1 7551 13 0.23 0.08121 0.00113 0.06803 0.00217 0.06653 0.00142 17.9 0.4

6.1 7848 10 0.28 0.08330 0.00142 0.06459 0.00215 0.06459 0.00215 17.2 0.4

7.1 9457 14 0.28 0.05235 0.00065 0.09992 0.00292 0.09992 0.00292 14.6 0.3

8.1 13328 9 0.19 0.07181 0.00087 0.06193 0.00132 0.06193 0.00132 17.0 0.3

9.1 7730 14 0.28 0.05907 0.00075 0.06593 0.00293 0.06593 0.00293 15.1 0.3

10.1 9430 13 0.24 0.05449 0.00076 0.06550 0.00227 0.06550 0.00227 14.6 0.3

11.1 10681 16 0.14 0.07113 0.00057 0.06687 0.00172 0.06687 0.00172 16.6 0.2

12.1 10751 12 0.34 0.07715 0.00145 0.07498 0.00316 0.07498 0.00316 17.3 0.4

13.1 11880 11 0.50 0.07158 0.00093 0.14395 0.00400 0.14395 0.00400 16.2 0.3

14.1 14015 13 0.16 0.06976 0.00118 0.06222 0.00156 0.06222 0.00156 16.8 0.4

15.1 13242 10 0.16 0.07083 0.00139 0.06123 0.00121 0.06123 0.00121 16.7 0.4

E.2.6 NX9439 Naxos S-type Granite (Z2037, 97795)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 208Pb /

232Th

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

1.1 41087 4 3.18 0.02526 0.00042 0.06485 0.00306 0.06485 0.00306 12.1 0.1

2.1 43246 6 3.08 0.01871 0.00033 0.07073 0.00321 0.07073 0.00321 11.6 0.2

3.1 18543 21 1.87 0.02227 0.00042 0.09271 0.00426 0.09271 0.00426 12.3 0.2

3.2 18373 12 10.50 0.02709 0.00065 0.18623 0.00839 0.18623 0.00839 11.7 0.2

4.1 14350 21 2.03 0.02573 0.00050 0.09386 0.00445 0.09386 0.00445 12.4 0.2

5.1 16500 17 3.39 0.02259 0.00046 0.11059 0.00538 0.11059 0.00538 11.1 0.2

5.2 19504 14 8.83 0.02653 0.00102 0.17893 0.01952 0.17893 0.01952 11.4 0.2

5.3 23060 21 1.59 0.01902 0.00037 0.08448 0.00707 0.08448 0.00707 11.8 0.4

5.4 21455 10 1.70 0.02099 0.00067 0.06644 0.00324 0.06644 0.00324 11.9 0.2

6.1 27717 6 5.85 0.02517 0.00069 0.08890 0.00403 0.08890 0.00403 11.8 0.2

6.2 19076 11 30.45 0.04335 0.00165 0.39526 0.02625 0.39526 0.02625 12.2 0.3

6.3 15038 12 27.03 0.03857 0.00082 0.36858 0.01077 0.36858 0.01077 11.5 0.2

Appendix E 315

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 208Pb /

232Th

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

7.1 18846 22 1.56 0.02278 0.00039 0.08534 0.00603 0.08535 0.00603 12.3 0.2

7.2 13339 15 2.63 0.02653 0.00057 0.09234 0.00627 0.09234 0.00627 11.8 0.2

8.1 15509 16 2.24 0.02483 0.000621 0.09100 0.00345 0.09100 0.00345 12.4 0.2

9.1 10793 24 2.27 0.02719 0.00063 0.10344 0.00536 0.10344 0.00536 11.7 0.2

10.1 9244 22 3.85 0.02494 0.00064 0.13849 0.01152 0.13849 0.01152 11.3 0.3

11.1 9462 23 3.25 0.02866 0.00097 0.12563 0.01875 0.12563 0.01875 12.1 0.2

11.2 11648 25 2.25 0.02463 0.00056 0.11054 0.01111 0.11054 0.01111 12.5 0.2


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 208Pb /

232Th

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

1.1 4555 36 0.10 0.06918 0.00146 0.09034 0.00260 0.09034 0.00260 15.2 0.5

2.1 3790 17 0.52 0.06548 0.00097 0.09723 0.00645 0.09723 0.00645 14.5 0.4

3.1 3191 35 0.26 0.07307 0.00114 0.10125 0.00647 0.10125 0.00647 15.1 0.3

4.1 3989 39 0.16 0.06410 0.00100 0.09597 0.00452 0.09597 0.00452 14.8 0.3

5.1 3952 34 0.22 0.06238 0.00093 0.09655 0.00505 0.09655 0.00505 14.3 0.3

6.1 4249 46 0.19 0.05086 0.00127 0.16738 0.00528 0.16738 0.00528 13.2 0.5

7.1 3276 39 0.27 0.06156 0.00189 0.10225 0.00725 0.10225 0.00725 13.8 0.6

8.1 4294 34 0.17 0.06499 0.00133 0.09055 0.00429 0.09055 0.00429 14.4 0.4

9.1 5868 33 0.15 0.06436 0.00110 0.07973 0.00338 0.07974 0.00338 14.4 0.3

10.1 2836 31 0.29 0.07905 0.00195 0.09945 0.00638 0.09945 0.00638 16.3 0.6

12.1 1512 42 0.57 0.07210 0.00263 0.14559 0.01679 0.14559 0.01679 16.8 0.9

13.1 1961 33 0.48 0.08218 0.00207 0.14214 0.01126 0.14214 0.01126 16.7 0.6


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 208Pb /

232Th

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

2.1 4140 48 0.18 0.05683 0.00202 0.08391 0.00575 0.08391 0.005751 12.6 0.7

3.1 21193 5 3.64 0.10732 0.00213 0.32722 0.01414 0.32722 0.014140 9.8 0.4

5.1 2933 73 0.14 0.06315 0.00199 0.10358 0.00634 0.10358 0.00634 12.1 0.5

6.1 2729 63 0.70 0.07750 0.00461 0.14190 0.02529 0.14190 0.02529 12.3 1.1

7.1 3678 51 1.28 0.06359 0.00527 0.13680 0.03889 0.13680 0.03889 11.8 1.4

9.1 3712 54 0.22 0.04579 0.00118 0.10852 0.00772 0.10852 0.00772 11.5 0.4


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 208Pb /

232Th

±1σσσσ 207Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

10.1 2906 66 1.36 0.17706 0.00470 0.60480 0.01444 0.60480 0.01444 9.2 0.4

11.1 2799 92 0.30 0.05918 0.00358 0.13299 0.01532 0.13299 0.01532 12.2 1.0

11.2 2721 96 0.25 0.05670 0.00202 0.12334 0.01352 0.12334 0.01352 12.0 0.6

12.1 2066 90 0.52 0.07644 0.00221 0.14169 0.02514 0.14169 0.02514 12.3 0.5

13.1 3728 56 0.29 0.06980 0.00250 0.10555 0.01047 0.10555 0.01047 12.5 0.7

14.1 5259 43 1.13 0.11590 0.00434 0.40829 0.01724 0.40829 0.01725 11.8 0.7

15.1 4340 49 0.54 0.06555 0.00193 0.10690 0.01640 0.10690 0.01640 11.8 0.5

Appendix E 317

E.3 Titanite U-Pb Analytical Results


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ 208Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

3.1 1830 0.62 74 0.15210 0.00305 0.62270 0.01481 1.55945 0.03477 14.6 0.4

4.1 789 0.19 83 0.26721 0.00619 0.69613 0.02168 1.67697 0.04427 16.2 0.5

5.1 777 0.12 85 0.25972 0.00910 0.71005 0.02350 1.78728 0.06052 12.7 0.5

6.1 1362 0.35 72 0.14841 0.00363 0.61151 0.01565 1.51059 0.04044 14.1 0.4

7.1 1588 0.84 73 0.17058 0.00358 0.61686 0.01268 1.54664 0.04062 15.2 0.4

8.1 1172 0.48 75 0.17810 0.00430 0.62607 0.01509 1.55626 0.03751 15.6 0.4

8.2 2955 0.44 58 0.09552 0.00332 0.49974 0.01797 1.25441 0.04123 15.0 0.6

9.1 722 0.66 84 0.27343 0.00819 0.70157 0.01988 1.73042 0.05436 15.1 0.5

10.1 656 0.12 86 0.32955 0.01192 0.71598 0.02639 1.79455 0.06803 16.0 0.6

11.1 824 0.19 89 0.26445 0.00967 0.74075 0.02677 1.78274 0.06310 9.8 0.4

12.1 1159 0.36 80 0.17582 0.00734 0.66576 0.02853 1.67418 0.07388 13.9 0.7

13.1 1070 0.34 81 0.23402 0.00693 0.67887 0.01853 1.72056 0.06806 15.4 0.5

14.1 584 0.40 90 0.42786 0.01981 0.74581 0.02499 1.86423 0.06464 14.5 0.7

15.1 1507 0.39 69 0.14573 0.00440 0.58500 0.02005 1.42215 0.04791 14.9 0.5

15.2 2139 0.34 70 0.14254 0.00367 0.59338 0.02131 1.46322 0.05250 17.0 0.7

16.1 1062 0.20 85 0.21440 0.00525 0.70524 0.01646 1.70131 0.04888 12.1 0.4

16.2 1220 0.30 75 0.19445 0.00449 0.63293 0.02346 1.68164 0.04146 16.3 0.4

17.1 1009 0.42 83 0.21392 0.00456 0.69380 0.02184 1.65828 0.04539 12.8 0.4

18.1 954 0.27 80 0.22538 0.00544 0.66891 0.01524 1.58915 0.04267 14.6 0.4

18.2 1587 0.34 78 0.19617 0.00548 0.65673 0.02251 1.66781 0.05246 17.4 0.7

19.1 1277 0.62 79 0.20786 0.00386 0.66181 0.01418 1.68911 0.04198 15.7 0.4

20.1 1595 0.26 74 0.12680 0.00348 0.62404 0.01942 1.49098 0.06626 11.5 0.4


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ 208Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

1.1 2175 0.05 87 0.20107 0.00849 0.73512 0.01865 1.53598 0.16395 15.0 1.1

2.1 1140 0.12 86 0.30372 0.00386 0.73385 0.00854 1.68287 0.04157 13.0 0.2

4.1 1157 0.14 79 0.21599 0.00397 0.67506 0.00921 1.60640 0.04917 13.8 0.3

5.1 999 0.12 86 0.26501 0.01295 0.73398 0.01002 1.62851 0.04245 13.5 0.8


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ 208Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

6.1 264 0.03 93 0.65838 0.03032 0.78722 0.00931 1.89518 0.02243 13.5 0.6

7.1 792 0.08 83 0.29557 0.00770 0.70732 0.01956 1.69471 0.03581 15.3 0.4

8.1 639 0.17 86 0.38049 0.02296 0.73474 0.01273 1.81418 0.02975 13.9 0.8

9.1 611 0.14 83 0.28012 0.00515 0.70441 0.00693 1.72465 0.02042 12.0 0.2

10.1 298 0.08 91 0.51393 0.00810 0.76778 0.01055 1.83341 0.05114 11.8 0.2

11.1 452 0.09 84 0.38422 0.00555 0.71345 0.00976 1.64878 0.06242 14.9 0.3

12.1 972 0.11 77 0.22354 0.00315 0.66235 0.00919 1.61563 0.02343 14.0 0.2

13.1 532 0.24 86 0.36074 0.00415 0.72761 0.00754 1.70853 0.05979 12.9 0.2

14.1 562 0.34 90 0.46640 0.00801 0.76293 0.01193 1.87623 0.03183 13.1 0.2

E.3.3 NX9435 Naxos Amphibolite (Z2265, 97799)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ 208Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

1.1 1626 0.67 46 0.10736 0.00653 0.41979 0.09475 1.07809 0.17695 11.5 0.7

2.1 885 0.08 71 0.15876 0.02765 0.62771 0.16359 1.47748 0.31351 9.3 1.6

3.1 263 0.71 78 0.25178 0.04370 0.68378 0.10005 1.78333 0.27202 9.9 1.7

4.1 720 0.12 67 0.19464 0.03330 0.58938 0.09439 1.41202 0.20861 13.5 2.3

5.1 1098 0.10 45 0.12625 0.01306 0.41509 0.07010 0.99959 0.13642 14.2 1.5

6.1 1617 0.60 48 0.13842 0.01816 0.44010 0.08412 1.05722 0.14095 13.4 1.8

7.1 421 0.15 80 0.31111 0.07073 0.70187 0.12434 1.68519 0.25761 12.1 2.7

8.1 1069 0.06 52 0.13451 0.01640 0.46873 0.06875 1.07088 0.17037 13.4 1.6

9.1 916 1.30 68 0.20860 0.03909 0.59828 0.06325 1.57390 0.17451 13.9 2.6

10.1 879 0.17 64 0.21624 0.04078 0.56968 0.11871 1.39476 0.29043 15.3 2.9

11.1 996 0.09 59 0.18170 0.04157 0.52789 0.13911 1.23404 0.29724 14.8 3.4

12.1 470 0.14 76 0.27915 0.09078 0.66208 0.20619 1.57536 0.44135 12.7 4.1

13.1 749 0.17 68 0.19867 0.03994 0.59729 0.11456 1.40933 0.22791 13.9 2.8

14.1 1490 1.71 48 0.13762 0.02354 0.43573 0.10363 1.34318 0.21568 14.7 2.5

15.1 452 0.22 75 0.25102 0.05550 0.66067 0.11824 1.58228 0.20565 12.1 2.7

15.3 271 0.44 92 0.69377 0.18505 0.79390 0.14398 1.83733 0.32840 11.6 3.1

16.1 480 0.17 77 0.26149 0.05431 0.67153 0.10993 1.53820 0.23420 11.8 2.4

17.1 535 0.37 72 0.25511 0.05378 0.63364 0.14161 1.46669 0.31584 14.2 3.0

18.1 1185 0.36 61 0.17184 0.02381 0.54283 0.06310 1.25857 0.14622 13.2 1.8

19.1 464 0.20 69 0.18084 0.02366 0.60679 0.08497 1.35907 0.14922 11.2 1.5

20.1 906 0.47 61 0.16568 0.03417 0.54468 0.12926 1.28276 0.26950 12.2 2.5

Appendix E 319

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ 208Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

21.1 435 0.12 73 0.26740 0.05152 0.64410 0.08790 1.62266 0.20046 13.7 2.6

22.1 972 0.04 68 0.21816 0.05324 0.59906 0.15040 1.36781 0.31708 14.4 3.5

23.1 546 0.13 78 0.28952 0.07978 0.68381 0.16009 1.62907 0.35353 13.7 3.8

24.1 748 0.21 78 0.29385 0.06268 0.68421 0.11885 1.60687 0.21273 13.9 3.0

E.3.4 NX9301 Naxos I-type Granodiorite (Z1858, Z2313, 97790)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ 208Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

1.1 965 1.05 68 0.18314 0.00602 0.57947 0.01572 1.55017 0.03545 10.8 0.4

3.1 614 1.40 71 0.21233 0.01278 0.60238 0.02726 1.62548 0.08067 11.4 0.7

3.2 895 1.06 72 0.21458 0.00671 0.60905 0.02168 1.53965 0.05494 11.2 0.4

3.3 752 0.49 73 0.21751 0.00544 0.61594 0.01344 1.49149 0.04257 11.4 0.3

4.1 773 1.18 85 0.29801 0.01676 0.71215 0.03097 1.77141 0.08040 9.1 0.5

6.1 556 1.00 80 0.23693 0.00473 0.66864 0.01786 1.67714 0.04772 8.7 0.2

6.2 613 0.86 80 0.26085 0.00947 0.66806 0.02469 1.69124 0.05330 10.0 0.4

7.1 645 0.47 83 0.27048 0.00726 0.69474 0.01782 1.69067 0.04509 8.9 0.2

8.1 576 1.31 78 0.26091 0.01107 0.65164 0.02077 1.64053 0.05161 11.0 0.5

8.2 661 1.65 78 0.24000 0.01033 0.65816 0.01877 1.72982 0.06110 10.2 0.4

8.3 1085 1.62 78 0.20531 0.00901 0.65692 0.02672 1.63718 0.06182 9.8 0.5

9.1 734 0.89 72 0.21695 0.00658 0.60987 0.02051 1.52116 0.04653 11.3 0.3

9.2 626 1.08 75 0.21196 0.00636 0.63139 0.01236 1.57131 0.03160 9.4 0.3

10.1 789 0.52 85 0.29670 0.00752 0.70972 0.02031 1.74821 0.04949 9.2 0.3

10.2 588 1.14 83 0.27994 0.01057 0.69299 0.02477 1.75276 0.05964 8.8 0.3

10.3 1200 0.50 81 0.21442 0.00919 0.67590 0.02418 1.59774 0.05439 9.2 0.4

10.4 765 1.03 78 0.23909 0.00573 0.65774 0.01709 1.65571 0.04732 10.4 0.3

10.5 963 0.53 80 0.25237 0.00515 0.66893 0.01221 1.63399 0.02653 10.6 0.3

11.1 559 0.79 80 0.28133 0.01115 0.67254 0.02225 1.68374 0.05234 9.9 0.4

12.1 344 0.24 88 0.39437 0.02416 0.73375 0.03746 1.84028 0.09587 9.3 0.6

13.1 875 0.41 83 0.24451 0.00736 0.69346 0.01516 1.68124 0.04465 9.8 0.4

13.2 398 0.38 84 0.33689 0.00735 0.70486 0.01513 1.72798 0.04108 10.2 0.2

1.1a 702 0.43 82 0.16292 0.00557 0.68985 0.01093 1.70207 0.03376 10.1 0.3

2.1a 583 0.43 78 0.15125 0.00299 0.65382 0.01135 1.65085 0.02464 11.3 0.2

3.1a 453 0.29 84 0.20324 0.00835 0.70387 0.02145 1.76768 0.04747 10.8 0.4

4.1a 747 0.45 77 0.13661 0.00300 0.65183 0.01007 1.59196 0.02117 10.7 0.2

5.1a 1190 0.38 87 0.22276 0.00837 0.72384 0.00944 1.79953 0.02020 11.0 0.4


Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ 208Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

1.1b 1060 0.65 70 0.13473 0.00181 0.61661 0.00653 1.53992 0.01802 11.2 0.2

2.1b 898 1.27 65 0.11144 0.00190 0.57338 0.00774 1.49972 0.02418 10.9 0.2

2.2b 992 1.30 63 0.11044 0.00146 0.55716 0.00809 1.49059 0.02556 11.9 0.2

3.1b 541 1.45 66 0.12428 0.00310 0.58039 0.01292 1.60955 0.03590 11.7 0.3

3.2b 769 1.39 68 0.11492 0.00278 0.60045 0.01234 1.59265 0.04148 11.0 0.3

3.2b 769 1.39 68 0.11492 0.00278 0.60045 0.01234 1.59265 0.04148 11.0 0.3

4.1b 719 0.94 67 0.13423 0.00170 0.59036 0.00903 1.51543 0.02517 12.0 0.2

4.2b 798 0.45 67 0.13162 0.00216 0.59131 0.00973 1.47811 0.02703 12.1 0.2

5.1b 404 0.86 68 0.15154 0.00330 0.59538 0.01545 1.54049 0.03479 11.6 0.3

6.1b 611 1.32 69 0.12491 0.00241 0.60265 0.01183 1.58681 0.03757 11.2 0.2

7.1b 567 1.16 67 0.13844 0.00323 0.58684 0.01573 1.56413 0.04180 13.0 0.3

7.2b 598 0.98 68 0.13702 0.00279 0.59887 0.01008 1.57484 0.03142 12.5 0.3

E.3.5 NX9303 Naxos Fractionated I-Type Granite (Z2313, 97791)

Spot U

(ppm)

Th/U f % 206Pb /

238U

±1σσσσ 207Pb /

206Pb

±1σσσσ 208Pb /

206Pb

±1σσσσ Age

(Ma)

±1σσσσ

1.1 544 1.25 65 0.13508 0.00393 0.57228 0.01373 1.54440 0.05811 12.5 0.4

2.1 567 1.50 64 0.13058 0.00307 0.56785 0.01177 1.54673 0.04109 12.2 0.3

3.1 654 1.80 65 0.12502 0.00301 0.57625 0.01345 1.59072 0.03396 12.3 0.3

4.1 617 1.47 65 0.13252 0.00316 0.57737 0.01327 1.54640 0.04567 12.6 0.3

5.1 769 2.01 66 0.12331 0.00268 0.58105 0.01751 1.61844 0.04007 11.5 0.3

6.1 744 1.30 71 0.16468 0.00454 0.62070 0.01023 1.62371 0.03573 13.4 0.4

7.1 679 1.70 70 0.13570 0.00503 0.61785 0.01190 1.69787 0.04354 11.2 0.4

8.1 882 0.95 74 0.15658 0.00223 0.64493 0.01202 1.62540 0.03076 11.4 0.2

9.1 670 2.14 65 0.13098 0.00245 0.57449 0.01192 1.60000 0.03213 12.4 0.2

10.1 676 1.21 67 0.12635 0.00282 0.59144 0.01426 1.55470 0.04243 11.5 0.3

11.1 976 0.59 61 0.10687 0.00238 0.54275 0.01245 1.37935 0.03101 11.2 0.3

Documents

Keay Thesis 1998