Paleoproterozoic orogenesis in the southeastern Gawler Craton, South Australia

Paleoproterozoic orogenesis in the southeastern GawlerCraton, South Australia*

A. REID1{, M. HAND2, E. JAGODZINSKI1, D. KELSEY2 AND N. PEARSON3

1Geological Survey Branch, Minerals and Energy Resources, PIRSA, GPO Box 1671, Adelaide, SA 5001, Australia.2Continental Evolution Research Group, Geology and Geophysics, University of Adelaide, SA 5005, Australia.3ARC National Key Centre for Geochemical Evolution and Metallogeny of Continents (GEMOC), Departmentof Earth and Planetary Sciences, Macquarie University, NSW 2109, Australia.

Integrated structural, metamorphic and geochronological data demonstrate the existence of acontractional orogen preserved in the ca 1850 Ma Donington Suite batholith along the eastern marginof the Gawler Craton, South Australia. The earliest structures are a pervasive gneissic foliationdeveloped in the Donington Suite and interleaved metasedimentary rocks. This has been overprintedby isoclinal and non-cylindrical folding, and zones of pervasive non-coaxial shear with north-directedtransport, suggesting that deformation was the result of orogenic contraction. SHRIMP U–Pb zircondata indicate that a syn-contractional granitic dyke was emplaced at 1846+ 4 Ma. Overprinting thecontractional structures are a series of discrete, migmatitic high-strain zones that show a normalgeometry with a component of oblique dextral shear. U – Pb zircon data from a weakly foliatedmicrogranite in one such shear zone give an emplacement age of 1843+ 5 Ma. Rare aluminousmetasedimentary rocks in the belt preserve a granulite-grade assemblage of garnetþbiotiteþplagioclaseþK-feldsparþ silicate melt that formed at *600 MPa and *7508C. Peak metamorphicgarnets are partially replaced by biotiteþ sillimaniteþ cordierite assemblages suggesting post-thermalpeak cooling and decompression, and are indicative of a clockwise P – T evolution. ChemicalU – Th –Pb electron microprobe ages from monazites in retrograde biotite yield a minimum estimate forthe timing of retrogression of ca 1830 Ma, indicating that decompression may be linked to thedevelopment of the broadly extensional shear zones and that the clockwise P – T path occurred duringa single tectonothermal cycle. We define this ca 1850 Ma phase of crustal evolution in the easternGawler Craton as the Cornian Orogeny.

KEY WORDS: Cornian Orogeny, Donington Suite, Gawler Craton, orogenesis, Paleoproterozoic.

INTRODUCTION

Paleoproterozoic magmatism at ca 1850 Ma occurs in a

number of Australian Proterozoic terranes (Wyborn

1988), including the Halls Creek Orogen (Page &

Hancock 1988), the Pine Creek Orogen (Needham et al.

1988) and the Mt Isa Inlier (Wyborn & Page 1983). The

Gawler Craton, the central component of the South

Australian Craton, also contains a significant magmatic

suite of this age (Parker et al. 1993). In each of these

terranes, 1850 Ma magmatism occurs at or near the

onset of prolonged magmatic, sedimentary and deforma-

tional histories (Page 1988; Myers et al. 1996). In the case

of the Gawler Craton, an 1850 Ma magmatic suite forms

the basement to Paleoproterozoic bimodal volcano-

sedimentary successions and Mesoproterozoic magma-

tism, both of which contain a variety of mineralisation

types and significant mineralisation potential (Figure 1)

(Daly et al. 1998). Understanding the evolution of the

basement to these sequences may provide clues as to the

tectonic framework within which these younger ther-

mal and sedimentary events occurred. This contribu-

tion is a study of the ca 1850 Ma orthogneisses of the

southeastern Gawler Craton.

GEOLOGICAL SETTING

The Gawler Craton preserves a record of Archean to

Mesoproterozoic continental evolution (Figures 1, 2)

(Drexel et al. 1993; Daly et al. 1998; Fanning et al. 2007).

The oldest units are Late Archean (2560 – 2500 Ma)

volcano-sedimentary rocks that were deformed during

the ca 2460 – 2440 Ma Sleafordian Orogeny (Daly et al.

1998; Swain et al. 2005). Intrusion of felsic magmas

occurred in the southern Gawler Craton at ca 2000 Ma

*Data Tables 1 and 2 [indicated by an asterisk (*) in the text and listed at the end of the paper] are Supplementary Papers; copies may

be obtained from the Geological Society of Australia’s website (5http://www.gsa.org.au4) or from the National Library of

Australia’s Pandora archive (5http://nla.gov.au/nla.arc-251944).{Corresponding author: [email protected]

Australian Journal of Earth Sciences (2008) 55, (449 – 471)

ISSN 0812-0099 print/ISSN 1440-0952 online � 2008 Geological Society of Australia

DOI: 10.1080/08120090801888594

Downloaded By: [University of Adelaide] At: 03:03 9 September 2008

(Miltalie Gneiss: Fanning et al. 1988), although little

is known about the tectonic setting of this event.

Following this, sedimentation of Hutchison Group

occurred, consisting of psammite and pelite with lesser

carbonate and iron-formation (Parker & Lemon 1982).

An intercalated volcanogenic unit within the upper

Hutchison Group (the Bosanquet Formation) has a

zircon U – Pb crystallisation age of 1866+ 10 Ma

(Fanning et al. 2007) and provides an upper bound on

the timing of sedimentation. It is noted that Vassallo &

Wilson (2001) have suggested that the Hutchison Group

may be younger than ca 1850 Ma and that the Bosanquet

Formation may represent a basement thrust slice that

has been tectonically juxtaposed with the Hutchison

Group. Recent detrital zircon investigations have re-

vealed a rock with a maximum depositional age ca

1790 Ma within units mapped as Hutchison Group

(Jagodzinski et al. 2006), although the majority of

detrital zircon investigations from the group yield

maximum depositional ages ca 2000 Ma (Fanning et al.

2007). The younger age may indicate infolding of a

younger sequence within the Hutchison Group during

the Kimban Orogeny. Nevertheless, the provenance and

stratigraphy of the Hutchison Group requires further

investigation.

Dominantly granitic and monzonitic rocks of

the Donington Suite were then emplaced at 1850 Ma

(Parker et al. 1993). Subsequent to this, the region

preserves a complex Late Paleoproterozoic to Early

Mesoproterozoic orogenic record with numerous

cycles of sedimentation, magmatism, metamorphism

and deformation lasting until around 1500 Ma

when granites of the Spilsby Suite were emplaced

in the southern Gawler Craton (Figure 1) (Daly et al.

1998).

The Donington Suite (Schwarz 2003) is an intrusive

magmatic complex that occupies a north – south-

trending belt around 600 km long and up to 80 km wide

along the eastern margin of the Gawler Craton

(Figure 2). The Donington Suite consists of rocks that

range in composition from gabbro, gabbronorite, char-

nockite, granodiorite to alkali granite and includes

within it the rocks of the Colbert Granite, formerly

known as the Colbert Suite (Mortimer et al. 1988a;

Hoek & Schaefer 1998; Schwarz 2003). Early workers

considered the Donington Suite to belong to the Lincoln

Complex, a rock association that was originally pro-

posed to encompass magmatism that occurred synchro-

nous with the Kimban Orogeny (Parker et al. 1993).

However, recent revision of this nomenclature (Schwarz

2003) has seen the Donington Suite excluded from the

Lincoln Complex as the timing and duration of the

Figure 1 Temporal evolution of the southeastern Gawler

Craton, showing magmatic, volcano-sedimentary, deforma-

tion/metamorphism and mineralisation events. For details

on individual events and rock units see Drexel et al. (1993),

Daly et al. (1998) and Fanning et al. (2007). Note that the

1850 Ma event is termed the ‘Cornian Orogeny’ rather than

previous descriptors (see Discussion). Ages quoted are

U – Pb zircon determinations from Fanning et al. (1988,

2007). Mineralisation is indicated as elemental occurrences

rather than deposit/prospect styles for simplicity: for

details on mineralisation styles in the Gawler Craton, see

Daly et al. (1998 and references therein).

3

450 A. Reid et al.


Figure 2 Geology of the Gawler Craton, South Australia. (a) Interpreted solid geology of the Gawler Craton, highlighting the

Sleaford Complex, Miltalie Gneiss, Hutchison Group, Donington Suite together with the widespread Gawler Range Volcanics

and Hiltaba Suite; the remaining Paleoproterozoic to Mesoproterozoic units are not differentiated and shown in grey. Major

mines are highlighted. (b) Outcrop map showing the locations of Donington Suite and other Proterozoic stratigraphic units of

the southeastern Gawler Craton (after Drexel et al. 1993).

Paleoproterozoic orogenesis, Gawler Craton 451


Kimban Orogeny have been clarified (see below).

Previous U – Pb geochronology of the Donington Suite

indicates that magmatic crystallisation of a range of

lithologies—including a quartz gabbronorite gneiss

used as a SHRIMP zircon age standard (Black et al.

2003)—occurred within error of 1850 Ma (Creaser &

Cooper 1993; Parker et al. 1993; Jagodzinski 2005).

Geochemically, Donington Suite granitoids show LREE

enrichment, negative Nb, Sr, P and Ti anomalies and

have eNd1850 Ma values between 72 and 74 (Schaefer

1998). The Donington Suite is thought to be derived from

a mixture of melt from a mafic parent—possibly a mafic

underplate—and crustal material (Mortimer et al. 1988a;

Schaefer 1998).

The major Paleoproterozoic orogenic phase recog-

nised in the southeastern Gawler Craton is the

Kimban Orogeny. Many workers considered the Kim-

ban Orogeny to have been long lived, occurring

between 1850 and 1700 Ma (Thompson 1969; Glen et al.

1977; Daly et al. 1998; Zang & Fanning 2001). However,

recently Hoek & Schaefer (1998) and Vassallo & Wilson

(1999, 2002) have shown that a tectonic foliation

developed within the Donington Suite prior to

1730 – 1700 Ma reworking. Mortimer et al. (1988a) also

identified a foliation in the Donington Suite that

developed before the emplacement of the Colbert

Granite, which also indicated a pre-1730 Ma timing

for this foliation. From these observations, Hoek &

Schaefer (1998) suggested that the earlier foliation

indicated the occurrence of a separate tectonothermal

event or orogeny, and that the latter, 1730 – 1700 Ma

event alone should be considered as the Kimban

Orogeny. This notion is supported by the development

of several phases of sedimentation and volcanism

between 1850 Ma and 1740 Ma (Figure 1) in the region,

interpreted by earlier workers (Parker et al. 1993; Daly

et al. 1998) to record the effects of the Kimban Orogeny.

Subsequently, the importance of this earlier ca 1850 Ma

event has been recognised and has been variously

termed the ‘Lincoln Orogeny’ (Vassallo & Wilson 1999)

or the ‘Neill Event’ (Ferris et al. 2002). Despite this

recognition, and the fact that the Donington Suite

dominates the crustal architecture along the eastern

margin of the Gawler Craton, little work has focused

on evaluating the structural and metamorphic expres-

sion of 1850 Ma tectonism.

One of the major difficulties in evaluating the

significance and character of the 1850 Ma event in the

eastern Gawler Craton is the intensity of Kimban

reworking of Archean and Paleoproterozoic rocks on

Eyre Peninsula (Parker 1980; Parker et al. 1993;

Vassallo & Wilson 2002; Tong et al. 2004). Fortunately,

Donington Suite granitoids are also exposed along the

southwestern coast of Yorke Peninsula, some 80 km east

of the main zone of known Kimban-aged deformation

(Figure 2). Confirmation that the earlier ca 1850 Ma

phase of tectonism is preserved here was provided by

Zang & Fanning (2001), who obtained a SHRIMP zircon

U – Pb metamorphic age of 1845+ 7.3 Ma from a gar-

net – biotite – sillimanite granulite facies paragneiss in-

tercalated with the Donington Suite. However, aside

from this age determination, the extent to which the

Yorke Peninsula gneisses as a whole preserve the effects

of an 1850 Ma tectonothermal event has not been

previously evaluated.

PALEOPROTEROZOIC GEOLOGY OF SOUTHERNYORKE PENINSULA

Rock types

On Yorke Peninsula, Paleoproterozoic basement is ex-

posed only on coastal platforms due to a thick Cambrian

to Holocene sedimentary cap (Figure 3) (Zang 2006).

Metasedimentary rocks of the Corny Point Paragneiss

(Zang & Fanning 2001) make up less than *5% of the

outcropping Paleoproterozoic rocks on southwestern

Yorke Peninsula. The principal exposure of metased-

imentary rocks is at Corny Point, where garnetiferous

cordierite-bearing, migmatitic quartzo-feldspathic

gneiss is the dominant lithology. A discussion of the

metamorphic evolution of these critical outcrops is

presented below.

The dominant Proterozoic rock types on southwes-

tern Yorke Peninsula, the Donington Suite (Figure 3),

are composed of two magmatic units: the Gleesons

Landing Granite and the Royston Granite (Zang 2002,

2006). SHRIMP U – Pb zircon geochronology indicates

that the Gleesons Landing Granite was emplaced at

1850+ 5 Ma (Zang 2006), 1850+ 3 Ma and 1850+ 2 Ma

(Jagodzinski et al. 2006). Two samples of the Royston

Granite yielded ages of 1850+ 12 Ma and 1849+ 11 Ma

(Zang 2006), which are essentially identical to the

emplacement age for the Gleesons Landing Granite.

The Gleesons Landing Granite is volumetrically domi-

nant and contains a suite of rock types including

syenogranite, adamellite, granodiorite and augen

orthogneiss that is everywhere foliated and migmatised.

Recrystallised, foliated and rafted mafic dyke remnants

within the Gleesons Landing Granite show back veining

from the enclosing felsic lithologies, suggesting they

were emplaced soon after or during the crystallisation of

the felsic host. We refer to these dykes as Type 1 dykes.

The Royston Granite has an adamellite to syeno-

granite composition (Zang 2002) and intrudes the

Gleesons Landing Granite as a series of discrete dykes.

The intrusion chronology of the Royston Granite

relative to the structural elements in the Gleesons

Landing Granite is particularly important in unravel-

ling the deformation history of the Donington Suite and

is discussed in detail below.

Intruding both the Gleesons Landing Granite

and Royston Granite are a suite of mafic dykes

that are generally straight-sided and variably recry-

stallised, which we refer to as Type 2 dykes. These

dykes cross-cut or strike subparallel to the foliation of

the host-rock. In places, such dykes are cross-cut by

pegmatite dykes, which appear on field criteria to be

related to the Royston Granite. However, elsewhere

cross-cutting relationships between straight-sided

mafic dykes and Royston Granite equivalents are not

observed, and it is possible that multiple generations of

younger dykes are present, as has been documented in

the Donington Suite on Eyre Peninsula (Mortimer et al.

1988b).

452 A. Reid et al.


Whole-rock geochemistry and Hf isotopiccomposition

We have investigated the whole-rock geochemistry

and zircon Hf isotopic composition of the Gleesons

Landing Granite and Royston Granite. Whole-rock

samples from a range of rock types were analysed for

major- and trace-element element composition at a

commercial geochemical operation, Amdel Laboratories

in Adelaide (Data Table 1*). The Hf isotopic signature

of felsic units was obtained through laser ablation-

multicollector inductively coupled plasma-mass

spectrometry (ICPMS) of individual zircons using facil-

ities in the Geochemical Analysis Unit at the GEMOC

National Key Centre, Macquarie University, Sydney.

Analytical methods followed those of Knudsen et al.

(2001) and Griffin et al. (2002). From the selected ICPMS

trace, integrated 177Hf/176Hf, 176Lu/177Hf and 176Yb/177Hf

ratios were calculated from which an epsilon (e) Hf value

was derived based on the decay constant for Lu of

1.94610711 y71 (Patchett et al. 1981; Tatsumoto et al.

1981). Depleted mantle Hf model ages have been

calculated based on measured 176Hf/177Hf compared

with a model depleted mantle with a present-day176Hf/177Hf¼ 0.28325 and 176Lu/177Hf¼ 0.0384 (after Griffin

et al. 2002).

Samples from the Gleesons Landing Granite and

Royston Granite show a trend towards higher MgO,

Al2O3, TiO and CaO with lower silica content

(Figure 4a), consistent with fractionation of a single

magmatic suite. REE patterns for felsic members of the

Gleesons Landing Granite and Royston Granite show

light REE enrichment, similar to the trend expected for

average continental crust (Figure 4b). Both Type 1 and

Type 2 mafic dykes are broadly tholeiitic in composition

(Figure 4c). They also show LREE enrichment compared

with typical MORB (Figure 4d). These data may indicate

either a component of crustal contamination in these

magmas or that the lithospheric source was enriched in

these elements.

Zircons from the Gleesons Landing Granite have

eHf1850 Ma values that range from 74.0 to 5.3 (sample

R639091; Table 1), with a mean value of 0.7+ 1.4.

Depleted mantle model ages calculated for these zircons

are in the range 2.0 to 2.4 Ga (Figure 4e; Table 1). A

sample of the Royston Granite shows similar eHf1850 Ma

values, with a range from 71.7 to 5.8 (sample R639097:

Figure 3 Proterozoic rocks of southwestern Yorke Peninsula.

(a) Coastal outcrops of Proterozoic gneisses, showing the

Donington Suite and Hutchison Group equivalent metasedi-

ments along with regional trends of the S1//S2 and S4

fabrics. (b) Total magnetic intensity (TMI) image. (c) Solid

geology interpretation. Low magnetism is shown by the

metasedimentary rocks of the Corny Point Paragneiss in

the north. The Gleesons Landing Granite, which occupies

the bulk of the area, can be differentiated into the more

highly magnetic region in the south, which corresponds to

layered gneiss and discordantly migmatised granodiorite

gneiss, and an area of less strongly differentiated magnetic

signal that is interpreted as megacrystic granite – gneiss.

Section marked X – Y shown in Figure 6.

3



Table 1) and a mean of 2.2+ 1.5. Depleted mantle model

ages calculated for zircons of this sample are identical to

that of the Gleesons Landing Granite sample, being in

the range 2.3 – 2.0 Ga (Figure 4e; Table 1). These data

suggest that the Donington Suite rocks were derived

from the fractionation of a mafic magma that incorpo-

Figure 4 Geochemical and isotopic composition of the Donington Suite on Yorke Peninsula. (a) Harker plot of selected

elements vs SiO2, to illustrate compositional trends in various members of the Gleesons Landing Granite and Royston

Granite. (b) REE plot for samples of the Gleesons Landing Granite and Royston Granite, normalised to chondrite

(Boynton 1984). Shown for comparison are the average continental crustal values of Weaver & Tarney (1984).

(c) AFM classification plot of Irvine & Baragar (1971) for the mafic rocks of the Donington Suite. (d) REE plot for samples

of the Type 1 and Type 2 mafic dykes, normalised to chondrite (Boynton 1984). Shown for comparison are REE values for

MORB from Sun & McDonough (1989). (e) Plot of 176Hf/177Hf vs age (Ma) showing location of Depleted Mantle curve (assuming

present-day 176Hf/177Hf¼ 0.28325 and 176Lu/177Hf¼ 0.0384 after Griffin et al. 2002). The slope of the line of best fit between the

analyses and the depleted mantle curve is determined by the assumed value of the crustal 176Lu/177Hf ratio of 0.015 (Patchett

et al. 1981).

454 A. Reid et al.


rated a component of pre-existing crustal material. The

crustal contaminant is likely to be material at least

2.3 Ga in age, as suggested by the Hf depleted mantle

model ages. These results are consistent with the

conclusions of previous, more detailed, geochemical

and isotopic investigations of the Donington Suite

(Mortimer et al. 1988a, b; Schaefer 1998) and further

strengthen the concept that the Donington Suite is a

geochemically homogeneous batholith over its consider-

able extent (Hoek & Schaefer 1998).

STRUCTURAL GEOMETRY, KINEMATIC ANDTEMPORAL EVOLUTION OF THE DONINGTONSUITE

Four phases of deformation, D1 to D4, can be distin-

guished in the Paleoproterozoic paragneisses and

orthogneisses on Yorke Peninsula. In this section, we

present integrated structural observations and U – Pb

zircon geochronological data from structurally con-

strained samples. Geochronological data were obtained

via SHRIMP II instruments at Curtin University, Perth,

and the Australian National University, Canberra.

Zircon separates were acquired through standard den-

sity and magnetic-separation techniques. Hand-picked

zircons were mounted into epoxy and imaged

under backscatter electron and cathodoluminescence

to determine internal structure of the grains. Details

of the SHRIMP analytical methods are given in

Appendix 1.

D1: gneissic fabric and leucosome formation

The first deformation event, D1, resulted in the forma-

tion of the gneissic fabric, S1, in both the Gleesons

Landing Granite and the Corny Point Paragneiss. This

fabric is defined by leucosomal segregations, and by

planar alignment of biotite or hornblende. Aside from

these fabric elements, no macroscopic D1 structures are

apparent.

D2: north-directed, non-coaxial compressionaldeformation

D2 structures in the Gleesons Landing Granite vary

from low-strain disruption of the S1 fabric through

foliation boudinage (Figure 5a), to higher strain intra-

folial, isoclinal (Figure 5b) and non-cylindrical folding

(Figure 5c) and zones of pervasive ductile shear

(Figure 5d). The occurrence of intrafolial isoclinal

folding suggests that the gneissic foliation in the

Gleesons Landing Granite is a composite S1//S2 fabric,

which, at the regional scale, strikes *2908 and has a

variable but dominantly southward dip (Figures 3, 6). D2

is associated with regionally significant zones of planar

Table 1 Summary of LAM-ICPMS Hf isotopic results.

Analysis Hf176/Hf177 1s Lu176/Hf177 Yb176/Hf177 U – Pb age Initial Hf eHf 1s TDM (Ga)

R639091 Gleesons Landing Granite [670258E, 6104230N (GDA 94 Zone 53)]

1 0.281533 0.000028 0.000734 0.017213 1850 0.281506 72.1 1.0 2.3

2 0.281721 0.000026 0.001183 0.029403 1850 0.281678 4.0 0.9 2.1

3 0.281531 0.000026 0.000561 0.013221 1850 0.281511 71.9 0.9 2.3

4 0.281690 0.000026 0.000956 0.022605 1850 0.281655 3.2 0.9 2.1

5 0.281626 0.000028 0.000781 0.018350 1850 0.281598 1.2 1.0 2.2

6 0.281676 0.000024 0.000901 0.022229 1850 0.281643 2.8 0.8 2.1

7 0.281615 0.000029 0.000399 0.009529 1850 0.281600 1.3 1.0 2.2

8 0.281720 0.000020 0.001007 0.024530 1850 0.281683 4.2 0.7 2.1

9 0.281476 0.000020 0.000638 0.016616 1850 0.281453 74.0 0.7 2.4

10 0.281801 0.000029 0.002397 0.065410 1850 0.281714 5.3 1.0 2.0

11 0.281592 0.000028 0.000732 0.016356 1850 0.281565 0.0 1.0 2.2

12 0.281568 0.000027 0.001437 0.039208 1850 0.281516 71.8 0.9 2.3

13 0.281582 0.000020 0.000427 0.010917 1850 0.281566 0.0 0.7 2.2

14 0.281559 0.000019 0.000595 0.015565 1850 0.281537 71.0 0.7 2.3

15 0.281633 0.000027 0.001270 0.032369 1850 0.281587 0.8 0.9 2.2

16 0.281622 0.000021 0.000587 0.015147 1850 0.281601 1.3 0.7 2.2

R639097 Royston Granite [668525E, 6104106N (GDA 94 Zone 53)]

1 0.281649 0.000026 0.000876 0.026460 1850 0.281617 1.8 0.9 2.2

2 0.281742 0.000024 0.000382 0.010211 1850 0.281728 5.8 0.8 2.0

2 0.281700 0.000024 0.000946 0.026292 1850 0.281666 3.6 0.8 2.1

3 0.281685 0.000016 0.000655 0.018533 1850 0.281661 3.4 0.6 2.1

3 0.281701 0.000027 0.000704 0.019741 1850 0.281675 3.9 0.9 2.1

4 0.281638 0.000026 0.000621 0.018214 1850 0.281615 1.8 0.9 2.2

4 0.281735 0.000026 0.000495 0.013818 1850 0.281717 5.4 0.9 2.0

5 0.281639 0.000033 0.001038 0.030415 1850 0.281601 1.3 1.2 2.2

5 0.281574 0.000029 0.000502 0.015189 1850 0.281556 70.3 1.0 2.2

6 0.281590 0.000026 0.001423 0.039062 1850 0.281538 71.0 0.9 2.3

6 0.281623 0.000020 0.000780 0.022871 1850 0.281595 1.0 0.7 2.2

7 0.281556 0.000022 0.001125 0.032726 1850 0.281515 71.8 0.8 2.3



ductile deformation as evidenced by a *10 km wide

zone of augen orthogneiss (Figures 3, 6). D2 kinematic

indicators including s-type porphyroclasts, C – C0 shear

fabrics (Figure 5d), and asymmetric folding show

consistent north vergence (Figure 6b). Thus, D2

is characterised by non-coaxial, north-directed

Figure 5 Structures observed in the Donington Suite. (a) Discordantly migmatised granodiorite gneiss of the Gleesons

Landing Granite, deformed by foliation boudinage at Royston Head. Notebook 19 cm long. Location 668100E, 6104140N. (b)

Layered granite – gneiss of the Gleesons Landing Granite, Point Yorke. Pen is 14 cm long. Location 699300E, 6100250N. Inset

shows F2 isoclinal folding, which is only present in Unit A of the Gleesons Landing Granite. Width of view of inset *10 cm.

(c). Down-plunge view of F2 non-cylindrical fold within layered granite – gneiss of the Gleesons Landing Granite, Point

Souttar. Pencil is 10 cm long. Location 709130E, 6135820N. (d) D2 shear fabric in megacrystic granite – gneiss of the Gleesons

Landing Granite, Berry Bay. This regionally significant shear fabric shows north-directed kinematics and implies D2

resulted from compressional deformation. Photograph taken looking west. Pencil is 10 cm long. Location 683300E, 6133580N.

(e) Dyke of feldspar-rich megacrystic gneiss of the Royston Granite, which cross-cuts the S2 foliation within surrounding

discordantly migmatised granodiorite gneiss, Royston Head. Pen is 12 cm long. Location: 669075E, 6104080N. (f) Example of

narrow D4 mylonite zone reworking the S2 fabric in the Gleesons Landing Granite, Royston Head. Pencil is 12 cm long.

Location: 668020E, 6103920N. All locations are given in GDA 1994 Zone 53 coordinates.

456 A. Reid et al.


Figure 6 Structural observations across southwestern Yorke Peninsula. (a) Outcrop map of Proterozoic rocks on

southwestern Yorke Peninsula with stereographic projections describing the principal structural elements of the region.

All stereonets are lower hemisphere, equal-area projections. (b) Schematic cross-section along X – Y.



contractional deformation and significant strain parti-

tioning. Importantly, D2 does not affect the Royston

Granite.

D3: north-directed contraction

D3 resulted in the formation of tight to open F3 folds

in all units of the Gleesons Landing Granite and

interleaved metasedimentary rocks. F3 folds generally

plunge shallowly to the west or east. Granitic dykes

of the Royston Granite cross-cut the D2 structures,

but are weakly foliated (Figure 5e), and in places

appear to intrude along the axial plane of F3 folds in

the Gleesons Landing Granite. As these dykes are

foliated, this implies they were emplaced post-D2 to

syn-D3.

A sample of K-feldspar-rich megacrystic granite dyke

of the Royston Granite (sample R639097) yielded euhe-

dral prismatic zircons, with generally blunt termina-

tions (Figure 7a). The grains display oscillatory zoning

typical of igneous crystallisation. Rare narrow, high-U

rims are observed, that in the majority of cases were too

thin to analyse, although a few grains with thicker

overgrowths were targeted. Twenty-five cores and

single-phase grains along with nine rims were analysed

(Table 2). All analyses of cores and single-phase grains

are concordant and yield a weighted mean 207Pb/206Pb

age of 1846+ 4 Ma (MSWD¼ 1.3; probability of fit¼ 0.12:

Figure 7b). Of the rim analyses, most are either within

error of the grain centres and/or show high errors or

high U. However, one rim (RH3.1), at 1771+ 9 Ma (1s), is

concordant and has a low Th/U.

We consider the 207Pb/206Pb age of 1846+ 4 Ma to be

the crystallisation age of this granite. The gneissic

fabric in the granite is interpreted to relate to meta-

morphism and deformation soon after intrusion as

evidenced by the analysis of seven zircon rims, which

yield a weighted mean 207Pb/206Pb age of 1846+ 12 Ma

(MSWD¼ 1.8; probability¼ 0.09) of near-identical age.

There is limited evidence (one grain) of zircon regrowth

or recrystallisation at ca 1770 Ma, although the signifi-

cance of this analysis (RH3.1) is uncertain.

D4: south-side-down extension with componentof dextral strike-slip

D4 deformation is manifest as a series of shear zones

that overprint all prior structural fabrics in all units of

the Gleesons Landing Granite. D4 shear zones vary from

discrete metre-scale, mylonitic shear zones (Figure 5f) to

zones of pervasive reworking of the S1//S2 fabric

with minimum widths in the order of tens of metres

(Figures 8, 9). Discrete D4 shear zones are defined by a

gneissic fabric, S4, and consistently show south-side-

down, normal kinematics and a shallow (158) to moder-

ately (608) west-plunging stretching lineation, L4

(Figures 6, 8). The geometry and stretching lineation

orientation of these shear zones suggest that D4 resulted

from extension coupled with a component of dextral

strike-slip deformation. The frequency of D4 shear zones

increases towards the south such that on southernmost

Yorke Peninsula, the S1//S2 fabric is almost completely

overprinted by S4.

At the regional scale, S4 dips moderately to the

southwest, although we note that zones of local north

dip are also observed (Figure 6). We interpret this as

synchronous folding of the S4 surface during the fabric

development, akin to the apparent folding of the lower

plate rocks observed in other extensional systems (e.g.

Norwegian Caledonides: Fossen 1992).

At a number of localities along the southernmost

Yorke Peninsula, pegmatite and microgranite dykes of

the Royston Granite are observed to truncate the

gneissic fabric of D4 shear zones, although in places

these microgranite dykes display a weak foliation,

subparallel to the shear zone boundaries (Figure 8).

These observations imply that these microgranites were

emplaced late-syn- to post-D4. Importantly also, Type 2

mafic dykes are restricted to those areas in which D4

deformation is most strongly expressed (Figure 6),

where they intrude subparallel to the S4 gneissic

foliation.

In order to constrain the timing of D4 deformation, we

have sampled one of the microgranite intrusions (sample

R698218, Figure 8). The majority of zircons from this

sample are euhedral to subrounded, show strong oscilla-

tory zoning, and are commonly mantled by metamict

rims up to 100 mm thick (Figure 7c). Twenty-nine

SHRIMP analyses of cores and whole grains with no

rims were collected from this sample (Table 2). Most

are concordant or near-concordant, and the dominant

population of near-concordant analyses produces a

weighted mean 207Pb/206Pb age of 1843+ 5 Ma (MSWD¼1.5; probability¼ 0.09; n¼ 16: Figure 7d).

Attempts were made to analyse rims that occur on

zircons from this population; however, these analyses

are strongly discordant, have very high U and common

Pb and do not yield reliable age information (Figure 7d;

Table 2). On the basis of their high U content, these

overgrowths may be related to late-stage residual

magmatic fluids. Thus, on the available evidence, the

late-syn-tectonic microgranite is considered to have

crystallised at 1843+ 5 Ma.

METAMORPHIC CONSTRAINTS

Paragneisses at Corny Point are one of the few localities

on Yorke Peninsula where diagnostic metamorphic

mineral assemblages occur. Lithologies at this locality

include garnet-bearing quartzo-feldspathic gneiss and

pods of calc-silicate (Figures 9, 10a, b). All of these

lithologies, except the calc-silicate, show complex net-

works of garnet-bearing leucosomes that both parallel

and cross-cut the S2 gneissic foliation (Figure 10c).

These leucosomes occur in rocks that contain a biotite-

defined foliation along with matrix of plagioclase and

quartz. This suggests that the leucosomes may have

formed via the general reaction: biþ sillþqzþplag¼gtþmelt+K-spar (Spear 1993). The peak assemblage

does not contain sillimanite, suggesting that this reac-

tion was terminated by the exhaustion of sillimanite.

In order to constrain the timing of leucosome

formation, we have dated zircons preserved in an S2-

parallel garnet-bearing leucosome in paragneiss units at

Corny Point (sample R698216, Figure 7). The sample

458 A. Reid et al.


contains abundant subhedral zircon, with oscillatory

zoning present in many grains (Figure 7e). Cores are

common, and are generally more translucent, with a

brighter cathodoluminescence. Thirty analyses yielded

generally concordant analyses that cluster around

1850 Ma along with a number of older ca 2200 – 1875 Ma

Figure 7 Results of SHRIMP zircon U – Pb geochronology. (a) Cathodoluminescence image of zircons from sample R639097. (b)

Concordia plot for sample R639097. (c) Cathodoluminescence image of zircons from sample R698218. Inset shows transmitted

light image of one typical zircon with metamict overgrowth. (d) Concordia plot for sample R698218. (e) Cathodoluminescence

image of zircons from sample R698216. (f) Concordia plot for sample R698216.



Table 2 Summary of SHRIMP U – Pb results.

Spot U

(ppm)

Th

(ppm)

232T/238U 206Pbc

(%)

206Pb*

(ppm)

206Pb*/238U 207Pb*/235U 207Pb*/206Pb* Disc.

(%)

Age (Ma)

+1s +1s +1s 207Pb/206Pb +1s

R639097 [668525E, 6104106N (GDA 94 Zone 53)]

301.1 108 78 0.7 0.04 30 0.3281 0.0031 5.0685 1.0579 0.1120 0.0005 0 1833 8

302.1 68 38 0.6 0.12 19 0.3294 0.0033 5.0340 1.2210 0.1108 0.0008 71 1813 13

303.1 107 78 0.8 0.03 30 0.3247 0.0031 5.0719 1.0666 0.1133 0.0005 2 1853 8

304.1 107 64 0.6 0.25 31 0.3339 0.0032 5.2012 1.1400 0.1130 0.0007 71 1848 11

305.1 233 114 0.5 0.00 66 0.3306 0.0030 5.1496 0.9469 0.1130 0.0003 0 1848 5

306.1 305 197 0.7 0.01 84 0.3208 0.0029 4.9900 0.9276 0.1128 0.0003 3 1845 5

307.1 117 87 0.8 0.04 32 0.3217 0.0031 4.9841 1.0691 0.1124 0.0005 2 1838 9

308.1 100 68 0.7 0.00 28 0.3298 0.0032 5.1442 1.1362 0.1131 0.0007 1 1850 11

309.1 151 98 0.7 0.01 43 0.3303 0.0031 5.1248 1.1051 0.1125 0.0007 0 1841 11

310.1 104 44 0.4 0.06 29 0.3267 0.0031 5.0735 1.0628 0.1126 0.0005 1 1842 9

311.1 82 36 0.4 0.00 23 0.3323 0.0033 5.0991 1.0971 0.1113 0.0006 72 1821 9

312.1 183 106 0.6 0.01 51 0.3274 0.0031 5.0927 1.0176 0.1128 0.0004 1 1846 6

313.1 131 60 0.5 0.16 37 0.3289 0.0035 5.1250 1.1795 0.1130 0.0006 1 1848 9

314.1 214 151 0.7 0.06 61 0.3305 0.0030 5.1390 1.0609 0.1128 0.0006 0 1844 10

315.1 105 79 0.8 0.05 30 0.3328 0.0032 5.1289 1.0759 0.1118 0.0005 71 1828 9

316.1 130 83 0.7 0.04 37 0.3316 0.0031 5.1783 1.0228 0.1133 0.0005 0 1852 7

317.1 130 69 0.6 0.02 37 0.3350 0.0031 5.2327 1.0190 0.1133 0.0005 71 1853 7

318.1 85 39 0.5 0.05 24 0.3273 0.0036 5.0265 1.2311 0.1114 0.0006 0 1822 10

319.1 117 72 0.6 0.04 33 0.3311 0.0035 5.1722 1.1496 0.1133 0.0005 1 1853 8

320.1 166 84 0.5 0.04 47 0.3291 0.0030 5.1368 0.9952 0.1132 0.0004 1 1851 7

321.1 162 93 0.6 0.05 46 0.3316 0.0036 5.1741 1.1660 0.1132 0.0005 0 1851 8

322.1 259 132 0.5 0.02 74 0.3329 0.0030 5.1853 0.9492 0.1130 0.0003 0 1848 5

323.1 171 112 0.7 0.01 48 0.3278 0.0030 5.0806 0.9754 0.1124 0.0004 1 1839 6

324.1 195 87 0.5 0.03 55 0.3299 0.0030 5.1219 0.9735 0.1126 0.0004 0 1842 6

325.1 171 73 0.4 0.01 49 0.3299 0.0030 5.1078 0.9772 0.1123 0.0004 0 1837 6

RH1.1 1009 376 0.4 0.10 300 0.3452 0.0058 5.3005 0.0937 0.1114 0.0006 75 1826 10

RH2.1 2164 633 0.3 0.03 628 0.3376 0.0053 5.2708 0.0853 0.1132 0.0004 71 1854 6

RH3.1 1226 174 0.1 0.08 331 0.3141 0.0052 4.6792 0.0804 0.1080 0.0006 0 1771 9

RH3.2 246 130 0.5 0.10 69 5.2004 0.1074 0.3261 0.0054 0.1157 0.0014 4 1890 22

RH4.1 625 781 1.3 8.14 107 0.1826 0.0036 2.6030 0.1764 0.1034 0.0067 56 1702 77

RH5.1 1140 144 0.1 0.12 342 0.3487 0.0058 5.4250 0.0943 0.1128 0.0006 74 1850 9

RH6.1 1376 131 0.1 0.00 375 0.3172 0.0053 4.8906 0.0861 0.1118 0.0006 3 1834 9

RH7.1 1803 314 0.2 0.05 511 0.3295 0.0027 5.1496 0.0516 0.1134 0.0007 1 1854 11

RH9.1 157 83 0.5 0.57 42 0.3123 0.0092 4.7150 0.2038 0.1095 0.0035 2 1791 58

R698218 [668871E, 6093551N (GDA 94 Zone 53)]

402.1 1592 1210 0.8 0.00 451 0.3298 0.0028 5.1420 0.0447 0.1131 0.0001 1 1849 2

404.1 129 97 0.8 0.05 37 0.3300 0.0032 5.1206 0.0538 0.1126 0.0005 0 1841 8

405.1 220 83 0.4 0.03 62 0.3275 0.0031 5.0980 0.0512 0.1129 0.0004 1 1846 6

406.1 180 96 0.5 0.54 51 0.3247 0.0035 4.9711 0.0620 0.1110 0.0007 0 1816 11

409.1 129 59 0.5 0.03 36 0.3259 0.0030 5.0578 0.0518 0.1126 0.0005 1 1841 8

410.1 231 107 0.5 0.01 66 0.3317 0.0030 5.2246 0.0494 0.1142 0.0003 1 1868 5

411.1 321 216 0.7 0.44 91 0.3298 0.0029 5.1142 0.0493 0.1125 0.0004 0 1840 7

413.1 408 117 0.3 1.38 93 0.2630 0.0024 3.3867 0.0585 0.0934 0.0014 71 1496 28

415.1 438 235 0.6 0.38 120 0.3176 0.0028 4.8628 0.0464 0.1110 0.0004 2 1816 7

416.1 233 96 0.4 0.09 65 0.3244 0.0031 5.0307 0.0538 0.1125 0.0006 2 1840 9

418.1 370 150 0.4 0.02 106 0.3330 0.0031 5.2180 0.0506 0.1136 0.0003 0 1858 4

419.1 195 69 0.4 0.03 55 0.3286 0.0030 5.1436 0.0504 0.1135 0.0004 1 1857 6

420.1 284 147 0.5 0.64 78 0.3186 0.0037 4.9035 0.0880 0.1116 0.0015 2 1826 25

421.1 372 250 0.7 0.02 105 0.3290 0.0029 5.1192 0.0470 0.1129 0.0003 1 1846 4

422.1 221 114 0.5 0.03 63 0.3339 0.0030 5.1779 0.0496 0.1125 0.0004 71 1840 6

423.1 232 145 0.6 0.06 65 0.3270 0.0030 5.0756 0.0489 0.1126 0.0004 1 1842 6

424.1 176 85 0.5 0.04 50 0.3279 0.0030 5.1050 0.0504 0.1129 0.0004 1 1847 7

425.1 286 122 0.4 0.06 79 0.3228 0.0029 4.9756 0.0468 0.1118 0.0003 1 1829 5

426.1 202 75 0.4 0.02 57 0.3298 0.0034 5.1016 0.0545 0.1122 0.0004 0 1835 6

427.1 352 117 0.3 0.78 86 0.2817 0.0096 4.3720 0.1816 0.1126 0.0027 15 1841 43

428.1 105 78 0.8 0.18 30 0.3326 0.0032 5.1321 0.0554 0.1119 0.0006 71 1831 9

429.1 298 149 0.5 0.25 83 0.3233 0.0030 5.0256 0.0497 0.1127 0.0004 2 1844 6

TG1.3 1778 289 0.2 3.04 393 0.2492 0.0022 3.6461 0.0569 0.1061 0.0014 21 1734 24

TG2.1 2639 184 0.1 7.76 529 0.2154 0.0027 2.6013 0.1050 0.0876 0.0034 9 1374 74

TG3.1r 11457 850 0.1 0.57 1436 0.1451 0.0010 1.4427 0.0128 0.0721 0.0004 13 989 12

TG3.2c 229 82 0.4 0.00 60 0.3061 0.0052 4.7353 0.1041 0.1122 0.0016 7 1836 25

TG4.1 2114 10731 5.2 40.79 610 0.1990 0.0087 1.8809 0.3186 0.0686 0.0112 724 886 338

(continued)

460 A. Reid et al.


ages that are interpreted to be detrital zircons sca-

venged from the surrounding metasedimentary rocks

(Table 2). A weighted mean 207Pb/206Pb age of

1848+ 8 Ma can be calculated from the 14 most con-

cordant analyses (MSWD¼ 1.4; probability¼ 0.14),

which is interpreted to record the timing of zircon

crystallisation in the leucosome. This age is within

error of the metamorphic zircon age reported by Zang &

Fanning (2001), thus confirming that high-temperature

metamorphism and leucosome formation in the Corny

Point Paragneiss is associated with the 1850 Ma event.

The peak garnet has been partially replaced by inter-

growths of biotiteþ cordierite+ sillimanite (Figure 10d,

e, f). In many instances, cordierite coronae isolate

garnet from matrix quartz (Figure 10f), while in other

examples, biotite and cordierite form pseudomorphs of

garnet (Figure 10e, f). This retrograde assemblage is

similar to those imaged on numerous P – T pseudosec-

tions (Harley & Carrington 2001; White et al. 2001) and

recorded in field studies (Clarke & Powell 1991;

Norlander et al. 2002). In the cited examples, the pro-

gression from the upper amphibolite to granulite facies

garnet-bearing assemblages to retrograde, biotiteþcordierite+ sillimanite-bearing assemblages occurs in

response to high-temperature decompression, and

in some instances, workers have inferred that

decompression was essentially isothermal (Clarke &

Powell 1991).

In order to evaluate the conditions of metamorphism

in the Corny Point Paragneiss, a calculated P – T

pseudosection has been constructed in the model system

Na2O – CaO – K2O – FeO – MgO – Al2O3 – SiO2 – H2O – TiO2 –

Fe2O3 (NCKFMASHTO; White et al. 2003) (Figure 11).

Mineral equilibria calculations were undertaken

using THERMOCALC 3.0 (Powell & Holland 1988) and

the internally consistent thermodynamic dataset of

Holland & Powell (1998). The bulk-rock composition

used in the calculations was taken from whole-rock

XRF analyses of a garnet – cordierite – biotite-bearing

metapelite from Corny Point (see Figure 11 for

composition).

The inferred biþ sill breakdown reaction is observed

to occur at conditions around 7308C and 500 – 700 MPa

(Figure 11). This reaction produces a garnetiferous

assemblage and occurs in the presence of silicate melt,

as is indicated by the extensive network of garnet-

bearing leucosomes at Corny Point. The progression to

cordierite-bearing assemblages due to the breakdown of

garnet observed in the Corny Point Paragneiss is

indicated on the pseudosection to occur due to decom-

pression to 400 – 500 MPa at elevated temperatures.

The locally evident complete breakdown of garnet

Table 2 (Continued)

Spot U

(ppm)

Th

(ppm)

232T/238U 206Pbc

(%)

206Pb*

(ppm)

206Pb*/238U 207Pb*/235U 207Pb*/206Pb* Disc.

(%)

Age (Ma)

+1s +1s +1s 207Pb/206Pb +1s

R698216 [684154E, 6136920N (GDA 94 Zone 53)]

201.1 92 69 0.8 0.01 32 0.4065 0.0068 7.7240 0.1437 0.1378 0.0011 0 2200 14

202.1 145 65 0.5 0.06 41 0.3306 0.0051 5.1413 0.0886 0.1128 0.0009 0 1845 14

203.2 168 178 1.1 0.01 48 0.3298 0.0050 5.1838 0.0878 0.1140 0.0009 1 1864 14

204.1 1025 391 0.4 0.01 299 0.3398 0.0044 5.3913 0.0713 0.1151 0.0003 0 1881 5

204.2 958 337 0.4 0.01 291 0.3538 0.0046 5.5846 0.0749 0.1145 0.0004 74 1872 6

205.1 127 46 0.4 0.16 36 0.3340 0.0053 5.2362 0.0951 0.1137 0.0010 0 1859 16

205.2 113 41 0.4 0.26 32 0.3286 0.0052 5.1210 0.0974 0.1130 0.0012 1 1849 19

206.1 568 297 0.5 0.01 167 0.3415 0.0045 5.6177 0.0772 0.1193 0.0004 3 1946 7

206.2 546 240 0.5 0.30 150 0.3200 0.0042 5.1560 0.0719 0.1169 0.0005 7 1909 8

207.1 342 242 0.7 0.01 95 0.3222 0.0044 5.0268 0.0735 0.1132 0.0006 3 1851 9

207.2 313 211 0.7 0.06 81 0.3028 0.0042 4.6655 0.0690 0.1118 0.0006 7 1828 10

208.1 193 241 1.3 0.02 56 0.3381 0.0049 5.2972 0.0851 0.1136 0.0008 71 1858 12

208.2 205 227 1.1 0.02 57 0.3207 0.0047 4.9352 0.0808 0.1116 0.0008 2 1826 13

209.1 206 70 0.4 0.04 59 0.3328 0.0049 5.1919 0.0835 0.1131 0.0008 0 1850 12

210.1 211 135 0.7 0.05 61 0.3376 0.0049 5.3251 0.0859 0.1144 0.0008 0 1870 13

211.1 276 174 0.7 0.02 88 0.3698 0.0052 6.3346 0.0947 0.1242 0.0006 71 2018 9

211.2 1118 150 0.1 0.54 293 0.3034 0.0047 4.7251 0.0823 0.1130 0.0009 8 1848 15

208.3 2201 15 0.0 0.06 508 0.2687 0.0038 4.0604 0.0601 0.1096 0.0005 17 1793 8

207.3 3210 115 0.0 0.08 586 0.2122 0.0030 2.9142 0.0422 0.0996 0.0004 30 1617 7

212.1 2847 68 0.0 0.55 568 0.2310 0.0033 3.2125 0.0490 0.1009 0.0006 22 1640 11

212.2 600 587 1.0 0.16 157 0.3031 0.0041 4.6859 0.0675 0.1121 0.0005 7 1834 9

214.1 446 44 0.1 0.07 136 0.3556 0.0051 5.8612 0.0884 0.1195 0.0006 71 1949 9

215.1 662 475 0.7 0.03 189 0.3320 0.0045 5.2000 0.0739 0.1136 0.0005 1 1858 8

215.2 1350 30 0.0 0.15 378 0.3258 0.0043 5.0347 0.0678 0.1121 0.0004 1 1834 6

216.1 348 165 0.5 0.02 108 0.3605 0.0052 6.0499 0.0926 0.1217 0.0007 0 1981 10

217.1 212 81 0.4 0.17 67 0.3655 0.0056 6.0269 0.1046 0.1196 0.0010 73 1950 14

218.1 576 709 1.3 0.00 169 0.3422 0.0047 5.3429 0.0769 0.1132 0.0005 72 1852 8

219.1 279 201 0.7 0.00 81 0.3364 0.0050 5.2055 0.0834 0.1122 0.0007 72 1836 11

220.1 133 85 0.7 0.15 44 0.3827 0.0068 6.4099 0.1269 0.1215 0.0011 75 1978 16

221.1 426 10 0.0 0.20 118 0.3209 0.0045 4.8885 0.0748 0.1105 0.0007 1 1807 11



Figure 8 Outcrop map of Donington Suite rocks at (a) Meteor Bay and (b) The Gap, southern Yorke Peninsula. Base map and

some structural data from Meteor Bay modified from Pedler (1976); otherwise all data from our observations. All stereonets

are upper hemisphere, equal-area projections.

462 A. Reid et al.


to biotiteþ cordierite suggests that the retrograde

path continued with further decompression and cooling

into the biotiteþ cordieriteþK-feldsparþplagioclase

(þliquidþquartzþ ilmenite) field, suggesting that pres-

sures may have reached as low as *300 MPa (Figure 11).

Thus, we suggest that the Corny Point Paragneiss

experienced a clockwise P – T evolution.

In order to constrain the timing of post-peak mineral

growth, we have dated monazite that is intergrown with

biotite around partially replaced garnet porphyroblasts

Figure 9 Outcrop map of Corny Point. Base map modified from Richardson (1978). Note that the southwestern outcrops at

Corny Point show zones of intense D4 ductile shear associated with a strongly downdip stretching lineation, L4. Kinematic

indicators including s- and d-type mantled porphyroclasts and C – C0 fabrics show south-side-down kinematics (see

photograph inset). We note, however, that local apparent reversals in kinematic shear sense are observed, which may be

associated with isoclinal folding of the S4 mylonitic layering during progressive south-side-down ductile shear. In addition,

although in many cases south-side-down kinematics can be shown, a number of shear zones lie parallel to the axial plane of

tight F3 folds and some lack convincing kinematic indicators. Thus, the shear sense across some of these shear zones is

uncertain, and it is possible that some may in fact have developed during D3. All stereonets are equal-area lower hemisphere

projections and mineral elongation lineations plot as dots, fold axes as crosses.



(Figure 12a). Analyses were undertaken using the

electron microprobe in situ chemical U – Th – Pb method

at the University of Adelaide using a CAMECA SX51

Electron Probe Micro Analyser following the methods of

Clark et al. (2005), Swain et al. (2005) and Rutherford

et al. (2006). Analysis of the Malagasy monazite (MAD,

514 Ma: Fitzsimons et al. 2005) during the analytical

session yielded a weighted mean age of 515+ 17 Ma

Figure 10 Metamorphic characteristics of metapelites in the Corny Point Paragneiss. (a) Garnet – biotite – quartzo-feldspathic

gneiss, with garnet-bearing leucosomes and leucocratic layering. Pen lid is 4 cm long. Location: 684105E, 6136965N. (b) Calc-

silicate pod in quartzo-feldspathic gneiss. Location 684095E, 6136980N. (c) Example of a garnet-bearing leucosome parallel to

lithological layering, similar to sample R698216 dated at 1848+ 8 Ma with SHRIMP U – Pb zircon. Location 684035E, 6136995N.

(d) Cordierite corona surrounding large garnet porphyroblast. Plane-polarised light; field of view 8 mm. Sample number

MB481-676. Location 683980E, 6136860N. (e) Cordieriteþbiotite pseudomorph of garnet. Minor sillimanite is also present in

this example. Plane-polarised light; field of view 12 mm. Sample number MBCP-6A. Location 683960E, 6136875N. (f) Biotite

corona around garnet. Plane polarised light; field of view 8 mm. Sample number MBCP03-9. Location 684100E, 6136970N. All

locations are given in GDA 1994 Zone 53 coordinates. Thin-sections stored at Geological Survey, PIRSA.

464 A. Reid et al.


(MSWD¼ 1.15; n¼ 48), giving confidence in the

accuracy of the age data for the Corny Point samples.

Results (Data Table 2*) show a predominance of ages ca

1830 Ma, with two samples (R674762 and R674765) giving

weighted mean ages of 1827+ 18 Ma and 1824+ 18 Ma

from 35 and 33 analytical points, respectively

(Figure 12a; Table 3). The monazite chemical data also

show a scatter of younger ages towards ca 1300 Ma,

although given that the technique cannot detect dis-

cordance, these younger ages probably reflect modifica-

tion of the mean population of ca 1830 Ma. Chemical

mapping of two monazite grains reveals a complex

internal structure, with irregular and patchy zonation

(Figure 12b), which may suggest these grains have

been hydrothermally altered (cf. Poitrasson et al. 1996)

and thus account for Pb loss and younger ages. The

age peak at ca 1830 Ma thus provides a minimum

constraint on the timing of retrograde biotite growth

and therefore the down-pressure evolution of the

paragneiss. Consequently, the retrograde evolution

is interpreted to reflect the waning effects of ca

1850 Ma orogenic processes rather than younger (e.g.

Kimban: 1730 – 1700 Ma) changes in thermobarometric

conditions.

DISCUSSION

1850 Ma orogenesis in the southeastern GawlerCraton: the Cornian Orogeny

The 1850 Ma event is characterised by the emplacement

of the Donington Suite into a compressional tectonic

environment within which contractional strain and

granulite-facies peak metamorphism were synchronous.

Contractional deformation was terminated by high-

temperature, extensional deformation and the emplace-

ment of Type 2 mafic dykes. Syn-extensional micro-

granite was emplaced at 1843+ 5 Ma, within error of

syn-contractional granite (1846+ 4 Ma), suggesting that

phases of contraction and extension occurred within a

short period of time, probably less than *10 million

Figure 11 P – T pseudosection for Corny Point Paragneiss. Bulk composition for these calculations from XRF data of Howard

(2006) and is indicated in the top left of the figure. Arrow indicates inferred P – T path for the Corny Point Paragneiss based on

mineral paragenesis described in the text.



years. Our geochronological constraints on the timing of

metamorphic decompression are also within error of

those for the extensional deformation, and we suggest

that decompression is likely to be directly related to the

onset of extension. Thus, our interpretation of the field

evidence and available geochronology is that contrac-

tional deformation was transient and that the entire

tectonothermal cycle occurred within *10 million years.

Detrital zircons analysed from the Corny Point

Paragneiss (Zang & Fanning 2001; Howard et al. 2006;

this study) yield ages as young as ca 1870 Ma, suggesting

that a maximum depositional age for the sedimentary

precursor may well approach the timing of emplace-

ment of the Donington Suite to within at most 20 million

years. These data suggest that the tectonic setting for

the 1850 Ma event is one that was characterised by an

Figure 12 Results of chemical U – Th – Pb electron microprobe monazite geochronology from Corny Point Paragneiss. (a)

Cumulative probability plot and histogram of ages derived from U – Th – Pb analyses, showing peak at ca 1830 Ma and

younger Pb loss. Inset shows setting of some of the monazites analysed (in centre of small radiation damage halos) within

retrograde biotite around a garnet porphyroblast. Plot of monazite results using AgeDisplay 2.25 (Sircombe 2004). (b)

Chemical maps of selected monazites for U, Th, Y and Pb. Field of view is 700 mm wide.

466 A. Reid et al.


ability to undergo rapid switches in tectonic mode,

such that sedimentation and granite emplacement

were followed by transient contractional deformation

before the onset of strike-slip to broadly extensional

deformation over a time interval in the order of 20

million years.

The above description of the 1850 Ma event shows it

to be far from the simple low-strain environment

envisioned previously (Mortimer et al. 1988a; Hoek &

Schaefer 1998) and necessitates a revision of models for

the Late Paleoproterozoic evolution of the southeastern

Gawler Craton. First, we suggest that the name for this

event be revised. Previously, this event was included as

an early part of the Kimban Orogeny (Drexel et al. 1993)

or has been termed the Lincoln Orogeny (Vassallo &

Wilson 1999) or Neill Event (Ferris et al. 2002). However,

these terms are unsatisfactory owing to their

derivation from localities on Eyre Peninsula where the

preserved structural and metamorphic record is domi-

nated by reworking associated with the 1730 – 1700 Ma

Kimban Orogeny. Further, the term ‘Lincoln’ is un-

satisfactory, since the term has already been used in the

definition of a rock suite (‘Lincoln Complex’). The term

‘Neill Event’ was proposed (Ferris et al. 2002) in the

absence of any structural, metamorphic or geochrono-

logical data to define the nature of this ‘event.’ We

believe the term ‘Neill Event’ is misleading, since the

locality from which the name is derived (Port Neill,

Eyre Peninsula) is the type locality for the Kalinjala

Mylonite Zone of Parker (1980), which equates to the

Kalinjala Shear Zone of Vassallo & Wilson (2002). Since

the Kalinjala Shear Zone is the best example of a crustal-

scale structure active during the 1730 – 1700 Ma Kimban

Orogeny, we consider naming the 1850 Ma events after

this locality to be unsatisfactory. Rather, we propose the

term Cornian Orogeny be used in place of previous

descriptors, in recognition of the excellent record of

1850 Ma tectonism that is preserved on western Yorke

Peninsula and in particular the exposures at Corny

Point.

Paleoproterozoic tectonics of the easternGawler Craton

Paleoproterozoic magmatic rocks across Australia, in-

cluding the Donington Suite, have often been considered

to have formed in response to melting of a mafic

underplate within an intracontinental environment,

since these rocks typically have I-type affinities, geo-

chemical attributes suggestive of shallow derivation in

high geothermal gradient regimes, and show intraplate

or equivocal affinities on tectonic discrimination dia-

grams (Etheridge et al. 1987; Wyborn et al. 1987;

Mortimer et al. 1988a; Wyborn 1988). Certainly, in the

case of the Donington Suite, there is no obvious

geochemical evidence to suggest formation in an active

margin setting (Mortimer et al. 1988a); consequently,

a model of intracontinental magmatism has been

advocated for the tectonic evolution of the region

(Parker et al. 1993).

However, it has also been suggested that the Doning-

ton Suite may have formed inboard of, or adjacent to, a

subduction zone on the basis of the geochemistry of

cross-cutting mafic dykes (Mortimer et al. 1988b) and the

physical similarities that the Donington Suite shares

with plate-margin magmas, including the co-existence of

felsic and mafic magmas (Hoek & Schaefer 1998). These

features may in fact point to the importance of plate

margin processes in the generation of the Donington

Suite.

In many Paleoproterozoic orogens, the evidence for

subduction-related magmatism and accretionary tec-

tonic processes is well established. Important examples

include the Trans-Hudson and Torngat Orogens of

North America (Scott 1998; St-Onge et al. 1999) and the

geodynamically related Nagssugtoqidian Orogen of

Greenland (van Gool et al. 1999) where juvenile crust

formation along magmatic arcs and the suturing of the

Archean Superior Province and the North Atlantic

Craton indicate that tectonic processes akin to Phaner-

ozoic plate tectonics were active at this time (van Gool

et al. 2002). Subduction-related tectonism and amalga-

mation of crustal blocks through the development of

wide orogenic belts have also been advocated for the

Paleoproterozoic evolution of Australia (Myers et al.

1996). Significant examples include the 1850 – 1830 Ma

Halls Creek Orogen, which records suturing of Kimber-

ley and North Australian Cratons (Bodorkos et al. 1999);

the 1830 – 1780 Ma Capricorn Orogen, which records the

suturing of the Pilbara and Yilgarn Cratons (Tyler &

Thorne 1990); and the broadly synchronous Paterson

Orogen, which resulted from the collision of the North

Australian and the West Australian Cratons (Bagas

2004).

Table 3 Summary of monazite chemical U – Th – Pb age data for samples from Corny Point.

Rock type n Range (ppm) Age (Ma) Age range MSWD

Pb Th U

R674762 [684063E, 6136976N (GDA 94 Zone 53)]

qz-felds-bi-cd-gt gneiss 35 3900 – 8290 313 900 – 56 500 3740 – 14550 1827+ 18 Ma 1.04

9 3880 – 6110 44 050 – 44 000 3460 – 10110 1640+ 53 to

1494+ 74 Ma

R674765 [684037E, 6136973N (GDA 94 Zone 53)]

gt-bearing leucosome 33 4250 – 8470 31 130 – 54 730 1410 – 12790 1824+ 18 Ma 0.8

35 2800 – 8330 1930 – 55 530 2200 – 13240 1312+ 67 to

1748+ 44 Ma



These examples suggest that plate-margin processes

were active during the Paleoproterozoic, and conse-

quently similar processes are highly likely to have been

involved in the formation of the Donington Suite. In the

absence of direct indications of subduction-related

magmatism, however, we suggest that the thermome-

chanical characteristics of the Cornian Orogeny itself

may provide clues as to a possible tectonic setting, in

particular the evidence for a clockwise P – T path, which

is indicative of collisional orogenesis, and the rapid

switch in tectonic mode between extension and

compression.

In Phanerozoic orogenic systems, changes in sub-

duction zone dynamics, such as the angle of the

downgoing slab or the arrival of a non-subductable

collider, play a decisive role in the deformational

history of regions above and inboard of the subduction

zone (Gutscher et al. 2000), and rapid switches from

extension to compression are common (Collins 2002a,

b). Likewise, regions hundreds of kilometres inboard

from an active subduction zone may also respond

rapidly to changes in far-field subduction dynamics, as

has been shown for the Paleozoic Lachlan Fold Belt of

eastern Australia (Collins 2002a) and invoked in recent

models for the evolution of Proterozoic Australia

(Giles et al. 2002, 2004).

On the basis of these considerations, we suggest a

possible tectonic scenario for the emplacement and

deformation of the Donington Suite as follows. To

reconcile the absence of subduction-related magma-

tism, we envisage that the Donington Suite was

generated in a (far-field?) continental backarc setting.

Lithospheric extension and consequent decompres-

sional melting of the mantle in the continental backarc

may have resulted in mantle-derived melts undergoing

both crustal assimilation and fractional crystallisation

as they ascended into the crust, thus forming the bulk

of the Donington Suite. The contractional deformation

that characterises the Cornian Orogeny may have been

focused into this backarc region due to thermal

softening induced by the thinning of the lithosphere

(Thompson et al. 2001) and the related emplacement of

the Donington Suite. Contractional deformation and

the clockwise P – T path may have been initiated as a

result of a change in the subduction dynamics, such as

the arrival of a buoyant collider, or a change in

subduction angle of the downgoing slab in the

hypothesised far-field subduction zone.

Following the contractional deformation, the rocks

appear to have undergone a phase of strike-slip to

obliquely extensional deformation, which is mimicked

in the metamorphic response by the decompressional

mineral assemblages. Decompression related to either

strike-slip or normal faulting is typical of metamorphic

core complexes of the North American Cordillera

(Norlander et al. 2002; Johnson 2006), and it is possible

that the switch from compression to oblique extension

in the later stages of the Cornian Orogeny records

similar unroofing processes. The driver for this decom-

pression may have been the renewal of subduction in

the outboard margin, triggering oblique extensional

deformation in the backarc.

In terms of the Late Paleoproterozoic evolution of the

Gawler Craton, the Cornian Orogeny represents only a

short interval (Figure 1). Subsequent to the Cornian

Orogeny, a suite of mafic dykes, the Tournefort Meta-

dolerite (formerly ‘Tournefort dykes’), were emplaced

into the Donington Suite (Schwarz 2003), and a number

of ca 1790 – 1740 Ma volcano-sedimentary sequences are

also preserved across the southeastern Gawler Craton

(Figure 1). The emplacement of the Tournefort Metado-

lerite at ca 1815 Ma (Schaefer 1998) and the subsequent

development of localised basins indicates extension is

likely to have continued periodically across the region

for up to 100 million years following the short-lived

convergence associated with the Cornian Orogeny.

Conceivably, this extension-dominated system over the

time interval 51850 to ca 1730 Ma may have developed

as a series of slab rollback-initiated, extensional basins

formed inboard of a far-field subduction front. This

hypothesised extension may have been terminated by

the onset of the Kimban Orogeny at ca 1730 Ma. This

type of tectonic setting may be likened to the extensional

accretionary orogen defined by Collins (2002a) in which

a long-lived subduction zone margin produces a series of

backarc basins that accrete to the continent through

successive, transient contractional orogenic events.

More work on the timing of sedimentation and deforma-

tion across the eastern Gawler Craton is necessary to

test this working hypothesis.

CONCLUSIONS

The ca 1850 Ma evolution of the southeastern Gawler

Craton was dominated by the emplacement of a large

granitic batholith, synchronous with high-grade con-

tractional deformation and terminated by high-

temperature decompression and extensional deforma-

tion. We define this magmatic and structural event as

the Cornian Orogeny. In terms of the Paleoproterozoic

evolution of the Gawler Craton, the Cornian Orogeny is

a short interval, and tectonic activity continued in the

region with the emplacement of mafic dykes and

bimodal volcano-sedimentary basin formation across

the southeastern Gawler Craton between ca 1810 Ma and

ca 1740 Ma, immediately prior to high-strain reworking

during the 1730 – 1700 Ma Kimban Orogeny. A working

hypothesis for the tectonic setting for the Paleoproter-

ozoic eastern Gawler Craton envisions a series of basins

forming inboard of a paleosubduction zone, possibly in a

continental backarc. Developing a clearer framework

within which the eastern Gawler Craton evolved during

the Paleoproterozoic will require further work on the

nature of the ca 1790 – 1740 Ma volcano-sedimentary

basins and integration of this with our understanding

of the 1850 Ma basement evolution.

ACKNOWLEDGEMENTS

This study is supported by Primary Industries

and Resources South Australia (PIRSA) via ARC

Linkage Grant LP0454301. Geoscience Australia is

468 A. Reid et al.


acknowledged for access to SHRIMP geochronology

facilities under the auspices of a National Geoscience

Agreement between Geoscience Australia and PIRSA.

Comments on earlier versions of this manuscript by

Alan Collins, Simon Bodorkos and Colin Conor along

with thorough reviews by Russel Korsch and Ron Berry

are acknowledged.

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Received 20 December 2006; accepted 27 October 2007

One to two kilogram samples of fresh rock were crushed

and heavy minerals isolated via a progressive washing

with a 10½’’ (35 cm) riffled gold pan. A 2 cm diameter

neodymium – iron – boron hand-magnet was then ap-

plied to remove magnetic minerals, followed by a

density separation in methylene iodide (3.3 g/mL).

Zircon selection was biased towards the least magnetic,

clearest grains, without discrimination between grain

morphologies. The zircon grains were mounted in

epoxy, together with the multi-grain zircon standard

QGNG and a small quantity of the standard SL13 (Black

et al. 2003). Grain mounts were polished to expose grains

in section, and all grains were subsequently photo-

graphed in transmitted and reflected light, and imaged

by cathodoluminescence on a Hitachi S2250 NSEM

located in the Electron Microscopy Unit at the Austra-

lian National University.

Analyses were made on the SHRIMP IIA ion

microprobe at Curtin University, Perth and the

SHRIMP II at Research School of Earth Sciences,

Australian National University, Canberra. A raster

time of 3 min was used, and data were acquired over

seven scans through the mass sequence. The primary

O27 beam was typically *4 – 5 nA, producing positive

secondary ions from elliptical spots 20 – 30 mm in size.

Differential fractionation between U and Pb was

monitored by reference to a 206Pb/238U ratio of 0.3341

for interspersed analyses of the 1850 Ma QGNG zircon

standard. Radiogenic Pb compositions were initially

determined by subtracting contemporaneous common

Pb (Stacey & Kramers 1975). All data were processed

using SQUID 1.12b (Ludwig 2001) and plotted using

ISOPLOT/EX 3.23 (Ludwig 2003). All weighted mean207Pb/206Pb ages determined from grouped data are

derived from 204Pb-corrected 207Pb/206Pb ratios. Ages

are calculated from the U and Th decay constants of

Jaffey et al. (1971), as recommended by Steiger & Jager

(1977).

SUPPLEMENTARY PAPERS

Data Table 1 Geochemical data.

Data Table 2 EPMA monazite data.

APPENDIX 1: SHRIMP GEOCHRONOLOGICAL METHODS



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Paleoproterozoic orogenesis in the southeastern Gawler Craton, South Australia