Buttneretal2005BasamentoQuilmes

Available online 3 June 2005

pegmatites. A younger group of pegmatites was emplaced in the country rocks at the end of the Ordovician (~440 Ma, RbSr

mineral isochrons), postdating most of the retrograde shear zones. Resetting of the muscovite KAr and 40Ar39Ar system in

Lithos 83 (2005) 143181

www.elsevier.com/locate/lithosAbstract

Towards unravelling the geodynamic setting of the northern Sierras Pampeanas (NW Argentina) we describe the tectono-

metamorphic and geochronologic evolution of sub-greenschist to granulite facies metamorphic sediments, granitoid plutons,

and pegmatites in the Ordovician Sierra de Quilmes metamorphic complex. The protoliths of the metasediments are represented

by a sequence of turbidites and minor calcsilicate rocks of the Neoproterozoic to Cambrian Puncoviscana Formation. The

metamorphic complex consists of four zones including the (1) chlorite, (2) biotitemuscovite, (3) garnetcordieritesillimanite,

and (4) orthopyroxene zones. Zones (3) and (4) show an increasing degree of anatexis, reaching large-scale diatexis in the

orthopyroxene zone at PT conditions exceeding ~800 8C and 600 MPa. At, or shortly after the metamorphic peak, the graniticto tonalitic Cafayate pluton intruded approximately along the boundary between anatectic and non-anatectic rocks. Retrograde

near-isobaric cooling of the middle crust was accompanied by non-penetrative ductile shearing at granulite to amphibolite facies

PT conditions. Evidence for significant prograde deformation is absent in the Sierra de Quilmes metamorphic complex.

Monazite and titanite UPb isotopic data constrain the metamorphic peak in migmatites and calcsilicate rocks to be at or

slightly prior to ~470 Ma. Retrograde amphibolite facies mineral reactions led to continuous formation of monazite and titanite

during slow cooling between ~470 Ma and 455 Ma (UPb data). The composite Cafayate pluton intruded over a time interval of

several million years between ~477 Ma (SmNd isochron) and ~460 Ma (monazite and titanite UPb isochron), followed byOrdovician metamorphism and plutonism in the Sierra de Quilmes

metamorphic complex: Implications for the tectonic setting of

the northern Sierras Pampeanas (NW Argentina)

S.H. Bqttnera,b,T, J. Glodnyc, F. Lucassenc, K. Wemmerd, S. Erdmannb,e,R. Handlerf, G. Franzb

aDepartment of Geology, Rhodes University Grahamstown, South AfricabInstitut fur Angewandte Geowissenschaften, Technische Universitat Berlin, Germany

cGeoForschungsZentrum Potsdam, GermanydInstitut fur Geologie und Dynamik der Lithosphare (IGDL), Universitat Gottingen, Germany

eDalhousie Department of Earth Sciences, Nova Scotia, CanadafInstitut fur Geologie, Universitat Salzburg, Austria

Received 30 April 2004; accepted 26 January 20050024-4937/$ - s

doi:10.1016/j.lit

T CorrespondiE-mail addree front matter D 2005 Published by Elsevier B.V.

hos.2005.01.006

ng author. Department of Geology, Rhodes University, Grahamstown 6140, South Africa.

ess: [email protected] (S.H. Bqttner).

covi

s me

f ma

phic

osed

lt tec

etting

dime

ronolo

LithosThe Ordovician high-grade metamorphism of the

Famatinian event was accompanied by the formation

of extensional marine sediment basins particularly at

the northern margin of the Sierras Pampeanas (e.g.

Bahlburg, 1991).

Despite a large number of publications, the

evolution and geodynamic setting of the Sierras

Pampeanas is still the subject of controversy. Two

conflicting models have been proposed, both with

internal variations by different authors and, in the light

of increasing data density and quality over the past 20

years, evolutionary modifications. One school of

thought follows concepts initially published by

Ramos et al. (1986) and Ramos (1988) that have

indicating subduction-related metamorphism are

absent. The concept of continent collision and terrane

amalgamation has been questioned in numerous

publications on the basis of field observation, bio-

stratigraphic, palaeomagnetic, sedimentological, geo-

chemical, geochronological, and petrological data

(e.g., Mon and Hongn, 1991; Bahlburg and Herve,

1997; Bock et al., 2000; Franz and Lucassen, 2001;

Acenolaza et al., 2002). The existence of proposed

terranes in the northern Sierras Pampeanas and

northern Chile has been challenged because of a

consistent and uniform geochronologic and petrologic

evolution across proposed suture zones (e.g., Lucas-

sen et al., 2000). In a contrasting model, Lucassen etweakly deformed pegmatites and crystallisation of new mus

attributed to minor late greenschist and sub-greenschist facie

The massive Early Ordovician heat transfer, the absence o

of non-penetrative deformation in the high-grade metamor

thickening and continent collision models, as have been prop

or stepwise extensional tectonics in a back-arc or a mobile be

the northern Sierras Pampeanas in general. An extensional s

agreement with the coeval formation of marine extensional se

Argentina and southern Bolivia.

D 2005 Published by Elsevier B.V.

Keywords: PTt(d) path; Anatexis; Near-isobaric cooling; Geoch

1. Introduction

The Sierras Pampeanas in northwestern Argentina

(Fig. 1) were consolidated at the southwestern margin

of Gondwana during the Middle Cambrian Pampean

and the Ordovician to Devonian/Silurian Famatinian

orogenies (e.g., Pankhurst and Rapela, 1998a; Lucas-

sen and Becchio, 2003). The Sierras Pampeanas

consist mainly of metamorphic equivalents of turbi-

ditic sediments of the late Neoproterozoic to Early

Cambrian Puncoviscana Formation, which have been

interpreted as passive margin deposits formed along

the palaeo-Pacific continental margin (e.g., Jezek et

al., 1985). Large parts of the Puncoviscana Formation

underwent regional metamorphism of up to granulite

facies conditions during both the Pampean and

Famatinian events, leading to large-scale anatexis

and widespread plutonism in the Sierras Pampeanas.

S.H. Buttner et al. /144favoured eastward subduction along an active Gond-

wana margin, leading to the formation of a magmaticte in low-grade metamorphic sediments at ~400416 Ma is

tamorphism and deformation.

jor prograde deformation, and subsequent, prolonged phases

zones during slow near-isobaric cooling contradict crustal

for the southern Sierras Pampeanas. We suggest continuous

tonic environment for the Ordovician Sierra de Quilmes, and

of the northern Sierras Pampeanas in the Ordovician is in

nt basins in vicinity of the Sierras Pampeanas in northwestern

gy; Extensional tectonics

arc and to successive and repeated collisional accre-

tion of exotic and parautochthonous terranes in the

latest Proterozoic and in the early Palaeozoic. These

models are based mainly on large-scale structural

patterns, tectonic models, and palaeo-magnetic data.

They have been widely accepted, used, and modified

in numerous later publications (e.g., Willner, 1990;

Dalla Salda et al., 1992; Ramos, 1995; Rapela et al.,

1998a; several contributions in Pankhurst and Rapela,

1998b; Omarini et al., 1999), or served as tectonic and

geodynamic background information for local and

petrogenetic studies undertaken in the Sierras Pam-

peanas (e.g., Otamendi et al., 1998, 1999).

However, typical indications for the subduction of

oceanic lithosphere, such as mafic or ultramafic rocks

that have been interpreted as ophiolites (Lottner and

Miller, 1986), are rare in the Sierras Pampeanas.

Blueschist or eclogite facies metamorphic rocks

83 (2005) 143181al. (2000) have interpreted the Pampean and Famati-

nian orogenies as contiguous stages in the evolution

ithosCalama

22 S

Very low-grade to unmetamor-phosed Puncoviscana FormationMetamorphic and igneous base-

Ordovician basin sediments

Boliv

ia70 W 66 W22 S

S.H. Buttner et al. / Lof an intra-cratonic mobile belt close to the south-

western Gondwana margin. The evolution of the

mobile belt culminated in widespread low-P/high-T

metamorphism at 525500 Ma, followed by a long-

standing high-thermal gradient regime in the mid

crust, which persisted until Silurian time (Lucassen

and Becchio, 2003).

The tectono-metamorphic evolution of the northern

Sierras Pampeanas is shown in a tilted segment of the

crustal migmatites and high-grade metamorphic

gneisses grade into upper-crustal medium- to very

Sierrade Quilmes

Tucuman

Salta

26 S

32 S70 W 66 W

Cafayate

? ment of the Sierras Pampeanas

Chile

Arge

ntin

a

200 km

26 S

32 S

Chile Argentina

Parag.

Ur.

Bra.

Fig. 1. Outcrop of the pre-Devonian basement (modified after

Lucassen et al., 2000; SEGEMAR, 1997) including the central and

northern Sierras Pampeanas, the Puncoviscana Formation, and the

exposed relicts of the Ordovician sediment basins in northwestern

Argentina. In the northern Sierra de Quilmes sediments of the

Neoproterozoic to Cambrian Puncoviscana Formation grade in

metamorphic rocks. Ordovician basins cover the basement of the

Sierras Pampeanas and the Puncoviscana Formation. Mesozoic and

Cenozoic westeast crustal shortening during the Andean Orogeny

caused the northsouth trending outcrop of all units. The box shows

the location of the map in Fig. 2.low-grade metasediments, exposing a metamorphic

complex in a tilted crustal segment (Fig. 2). We

consider the Sierra de Quilmes as a key area for the

understanding of the crustal evolution in the northern

Sierras Pampeanas, because rocks of various meta-

morphic grades are accessible which sheare the same

geological history at different levels of the crust.

The sedimentary protoliths of the metamorphic

rocks belong to the Neoproterozoic to Cambrian

Puncoviscana Formation, which consists mainly of

turbidites (Acenolaza et al., 1988; Jezek, 1990;

Omarini et al., 1999) deposited in a large sediment

basin that extended from Bolivia to central Argentina

(~338S) roughly between 648 and 688W (Rapela et al.,1990). Sediment sources were cratonic parts of south-upper to middle crust forming the Sierra de Quilmes

metamorphic complex. In this paper we present

integrated data on the metamorphic and geochrono-

logical evolution of a metapelitic to meta-greywacke

sequence in a high thermal gradient environment, and

correlate the PTt evolution with the timing of

regional tectonics and granitic and pegmatitic pluton-

ism. We discuss the tectono-metamorphic evolution in

the high-grade metamorphic basement in the context

of the coeval formation of sediment basins, which

have covered large parts of the Sierras Pampeanas in

northwestern Argentina. We propose a non-colli-

sional, most likely extensional, geodynamic setting

of the northern Sierras Pampeanas in the Ordovician

and Silurian periods.

2. Geologic setting of the Sierra de Quilmes

The Sierras Pampeanas form large northsouth

trending mountain ranges in central and northwestern

Argentina between 248 and 348S and 648 and 688W(Fig. 1) and consist mainly of gneisses, schists,

migmatites, phyllites, and, less commonly, marbles

and metabasites. Low-P/high-T metamorphism with

crustal anatexis, associated granitic plutonism and

polyphase deformation are typical and widespread

phenomena in the Sierras Pampeanas. At their north-

ern end, in the Sierra de Quilmes (~268S, 668W), mid-

83 (2005) 143181 145west Gondwana (Acenolaza et al., 1988; Schwartz and

Gromet, 2004). The maximum stratigraphic age is

Tolombn

Colalao del Valle

Cafayate

?

Chlorite zone

Biotite-muscovite zone

Garnet-cordierite-sillimanite zone

Cafayate pluton

Pegmatite swarms

Sample loc. for geochronology

Other loc. mentioned in text

Opx

Opx

Opx

Chl

Bt+Ms

532 Loc./sample No.

Orthopyroxene zone and unclas-sified high-grade metamorphic basement

PVCL/ sedimentary bedding

55 Cleavage and stretching lineation in ductile shear zones

52

?

?

?

?

?

Bt+Ms

Sil+CrdG

rt

?

?

?

?

?

43510 (K-Ar)

45011 (K-Ar)

45714 (K-Ar)

42012 (K-Ar)

468 (U-Pb)

4429 (Sm-Nd) (a)

4548 (Ar-Ar)

4087 (Ar-Ar)

4087 (Ar-Ar)

4428 (Ar-Ar)

4385 (Rb-Sr)

4405 (Rb-Sr)

441 (Rb-Sr)

47711 (Sm-Nd)4608 (U-Pb)

4603 (U-Pb)

~ 4028 (K-Ar)

473 (U-Pb)

~ 3858 (K-Ar)

459 (U-Pb)

552

MM 110MM72

540, 542

620

0 1 0km

MM 61

477

MM 22 MM 53

551

MM 26

(12 km south)

619

526523

515

513

465

557

482/83

488

496

510

472

506

532528

556537

554

Angastaco Formation (Miocene)

63

47

4651

67

56

78

52

56

69 58

43

6382

51

30

655762

73

5382

59

5649

71

55

45

55

6277

55

5174

48

62

76

32

39

51

49

65

San Antonio

S.H. Buttner et al. / Lithos 83 (2005) 143181146

unknown but is probably Neoproterozoic (Schwartz

and Gromet, 2004; and references therein). Sedimen-

tation in the Puncoviscana basin ended at ~530 Ma

(Durand, 1996, as cited in Rapela et al., 1998a).

of plutons within, and in the vicinity of, the study area

yielded Cambro-Ordovician ages (Rapela et al., 1982;

Miller et al., 1991). Lower Ordovician-aged plutonism

~50 km north of the study area has been established

(Becchio et al., 1999; Lucassen et al., 2001). Tectono-

metamorphic events of Mesozoic and Palaeogene age

ral n

ester

photo

S.H. Buttner et al. / Lithos 83 (2005) 143181 147Considerable proportions of the metamorphic rocks in

the Sierras Pampeanas, including the Sierra de

Quilmes, are metamorphic equivalents of Puncovis-

cana sedimentary rocks (e.g., Jezek et al., 1985; Rapela

et al., 1998a; Schwartz and Gromet, 2004).

The Middle Cambrian Pampean Orogeny is

believed to be a collisional event associated with

terrane accretion and late-orogenic extensional col-

lapse (Rapela et al., 1998a). Neither petrological nor

geochronological data suggest that the Pampean

Orogeny was of major importance in the Sierra de

Quilmes. Although Pampean deformation at low

temperatures of metamorphism might be hidden in

the metamorphic complex, the Pampean Orogeny is

not discussed in this study. During the consecutive

Famatinian Orogeny (~490350 Ma; e.g., Rapela et

al., 1998a) Ordovician sediments were deposited

discordantly on top of the Puncoviscana Formation

and other Cambrian sediments in marine extensional

basins (Bahlburg, 1991, 1998; Bahlburg and Herve,

1997; Bock et al., 2000). Relicts of these basins cover

large parts in northwestern Argentina and southern

Bolivia (Fig. 1), including the Puna, that adjoins the

Sierra de Quilmes on the western side. Apart from

minor and local occurrences of Early Silurian and

Early Devonian sediments, no post-Ordovician sedi-

ment record prior to the latest Carboniferous exists in

northwestern Argentina (Bahlburg and Herve, 1997).

Earlier studies on the Sierra de Quilmes have

presented petrographic and some geothermobaromet-

ric data from the southern part of the study area and

determined medium-pressure granulite facies meta-

morphism and deformation at high temperature (Rossi

de Toselli et al., 1976; Toselli et al., 1978). Rapela

(1976a,b) and Rapela et al. (1990) have studied the

geochemical composition of several plutons and

metamorphic Puncoviscana rocks in the northern

Sierra de Quilmes. RbSr WR (whole rock) dating

Fig. 2. The metamorphic complex of the Sierra de Quilmes. Mine

metamorphic zones (abbreviations after Kretz, 1983). Note that the w

Lithology and zone boundaries are assumed from aerial and satelliteof the metamorphic complex. (a)SmNd mineral isochron of a migmatitic

discussed in the text.are not documented, or are of minor importance

within the study area, and are therefore not described

here. Details concerning the post-Palaeozoic evolution

can be found in Del Papa (1999) and Salfity and

Marqillas (1994). The Palaeozoic basement is overlain

by fluvial conglomerates and arenites of the Miocene

Angastaco Formation (Hongn et al., 1998), in the

northern part of the study area (Fig. 2).

3. Petrography of metamorphic zones in the Sierra

de Quilmes metamorphic complex

Between San Antonio and Tolombon (Fig. 2), the

Sierra de Quilmes shows a continuous, northeast to

southwest transition from very low- and low-grade

metamorphic clastic sediments into amphibolite facies

and migmatitic rocks. The protoliths were mainly

turbiditic sediments of the Puncoviscana Formation

(e.g., Toselli, 1990; Toselli and Rossi de Toselli,

1990). Except for rare calcsilicate lenses and some

metabasites, this metamorphic zonation is defined in

rocks of a uniform psammo-pelitic to greywacke

ames along zone boundaries highlight the index mineral(s) of the

nmost parts of the area (marked with a b?Q) have not been mapped.graphs, gravel in river valleys, and extrapolation from mapped partsbased on UPb monazite data (467476 Ma; Lork and

Bahlburg, 1993). KAr WR ages betweem ~450 and

~470 Ma (Adams et al., 1990) from Puncoviscana

Formation metasediments in contact aureoles north of

the study area were interpreted as recording Ordovi-

cian contact metamorphism. More recent studies in

the area of Colalao del Valle (Fig. 2) indicated an age

of 442F9 Ma for the regional metamorphism anddeformation (SmNd mineral isochrons, Lucassen et

al., 2000). The radiogenic isotope composition (Sr,

Nd, Pb) of the Puncoviscana siliciclastic metasedi-

ments and the Cafayate pluton indicate large propor-

tions of recycled Palaeo- to Mesoproterozoic crustgneiss from Lucassen et al. (2000). The geochronological data are

Lithoscomposition. The composite granitic to tonalitic

Cafayate pluton crosscuts the metasediments. Follow-

ing the pluton emplacement, numerous pegmatites

penetrated the pluton and its country rocks.

The following sections describe important macro-

scopic and microscopic characteristics of the meta-

morphic zones and the plutonic rocks. In the high-

grade metamorphic zones biotite, sillimanite and

garnet have formed at different stages on the PT

path. Subsript b1Q indicates mineral phases that haveformed prior to or during partial melting, whereas

subscript b2Q refers to phases that have formed duringretrograde cooling. In cases, we also distinguish

between mineral assemblages in the neosome and

the mesosome. The latter consists of mineral assemb-

lages that have formed during prograde heating but do

not show evidence of anatexis. In the neosome

leucocratic minerals dominate, melanosomes are

generally rare. We use mineral reactions and petroge-

netic grids to determine minimum temperatures for the

peak and retrograde metamorphism in individual

metamorphic zones. Geothermobarometric data from

selected samples are presented in Section 4.

3.1. The chlorite zone

Very low-grade metamorphic Puncoviscana sedi-

ments are common to the east and north of San

Antonio (Fig. 2). This unit consists mainly of layered

psammo-pelitic to greywacke sedimentary rocks with

local occurrences of phyllites northeast of San

Antonio. Sedimentary structures including Bouma

sequences, compositional layering and cross-bedding

(Fig. 3a) are well preserved and dominate their

appearance in the field (Fig. 4a). Quartz-rich layers

are generally thicker than psammo-pelitic layers, the

average grain size is b10 Am, and there is no evidencefor temperature-controlled grain coarsening. Psammo-

pelitic layers consist of quartz, chlorite, illitic clay

minerals, extensively altered detrital biotite, white

mica, K-feldspar, and occasional plagioclase. Grey-

wacke layers are quartzfeldspar dominated. Locally,

a non-penetrative, steeply west-dipping crenulation

cleavage crosscuts the generally southeast dipping

sedimentary bedding (Fig. 2). Due to lack of

metamorphic assemblages suitable for thermobarom-

S.H. Buttner et al. /148etry, no precise PT data for the chlorite zone are

available. Illite crystallinity data from two samples(samples 620 A and C, Fig. 2) suggest very low-grade

metamorphic conditions (available in the Electronic

Data Supplement).

3.2. The biotitemuscovite zone

Increasing temperature of regional metamorphism

is indicated by the presence of muscovite and

brownish biotite neoblasts and by the breakdown of

chlorite suggesting the reaction:

Chl Kfs Ms Bt Qtz V 1(abbreviations after Kretz, 1983).

In the chemical system KFMASH, reaction (1)

takes place at upper greenschist facies conditions

(~440 8C; Simpson et al., 2000; Fig. 5a). Theassemblage of the biotitemuscovite zone is Qtz

BtMsPlFKfsFopaque mineral. Relicts of graded-and cross-bedding, and layers enriched in heavy

minerals indicate a sedimentary origin of the composi-

tional layering. Static quartz growth and recrystallisa-

tion become increasingly important towards the west

of the biotitemuscovite zone. The size of polygonal,

microscopically strain-free quartz grains in psammitic

layers increases from 30 to 50 Am near the chloritezone to 70150 Am at the margin to the garnetcordieritesillimanite zone, indicating temperature-

controlled grain coarsening. The meta-greywacke

layers do not show any secondary foliation. In the

psammo-pelitic layers, biotite and muscovite show a

shape-preferred orientation oblique to the sedimentary

bedding (Fig. 3b). The absence of cleavage domains

and microlithons, or any other evidence for pressure

solution, indicates low-strain conditions associated

with the generation of this foliation, or static oriented

mica nucleation in a non-hydrostatic stress field

(Passchier and Trouw, 1996).

The compositional layering of initially sedimentary

origin, manifested now by metamorphic or coarsened,

initially detrital minerals, is referred to as Puncovis-

cana compositional layering (PVCL, Figs. 3b and 4a).

Domains, layers, or rafts showing the PVCL are

present in all metasediments of the Cafayate area,

including the migmatites and diatexites in the high-

grade metamorphic zones, suggesting similar proto-

liths in all zones of the metamorphic complex.

83 (2005) 143181In the centre and the western part of the biotite

muscovite zone, microcline twinning in K-feldspar

1 mm

(a)

S2

PVCL

(b)

1 mm

Mes

BtNPlN

PlNQtzN

QtzN

(d)(c)

Sil1Bt1

700 m 1 mm

BtN

Opx

Crd

Pl

Grt

Grt

Pl

Bt1Grt

Pl

Qtz

(e) (f)

Crd/Pin

MsC

BtC

1 mm

Opx

BtC

MsRMsC

BtR

Opx3 mm

CrdSil+Ms

(g)

Opx1 mm

Mag QtzSil2

St

Crd

Crd

(h)

Opx50 m

S.H. Buttner et al. / Lithos 83 (2005) 143181 149

(j)

Lithos(i) CrdPin

Grt1

S.H. Buttner et al. /150suggests SiAl ordering during cooling below ~500 8C(Kroll et al., 1991). Further indication for amphibolite

facies metamorphism are chessboard subgrain patterns

in quartz, indicating PT conditions within the

stability field of high-quartz (Kruhl, 1996; Fig. 5a).

Bt2Bt1

Bt1

Grt2

Ky

500 m

CrdSil+Bt2

Bt1

Bt1

StKy Sil

StKy

Pin+Sil(k) (l)

350 m

Fig. 3. (a) Crossbedding in Puncoviscana Formation from the chlorite z

Sample 477B, plane-polarised light (PPL). (b) Static or low-strain foliati

(PVCL) obliquely. The absence of cleavage domains, microlithons, and elo

Sample 488A, PPL. (c) Peak metamorphic sillimanite1 and biotite1 from t

Migmatitic fabric in the garnetcordieritesillimanite zone. Quartz in n

subgrains and local grain boundary migration. Magmatic PlN is undeformed

growth during the formation of melt. The mesosome (Mes) consists of

polarised light (CPL). (e) Orthopyroxenegarnetcordierite assemblage of

Euhedral cotectic garnet shows flat zonation patterns (Fig. 7). The garnet

growth. Rim compositions are similar to those of garnet in sample 532A

Sample 557G, PPL. (f) Contact metamorphic growth of biotite (BtC), musc

strain regional foliation defined by muscovite (MsR) and biotite (BtR). Pin

contact aureole. Biotitemuscovite zone; sample 513, PPL. (g) Synkine

retrograde ductile shear zone. Sample MM22, CPL. (h) Formation of retrog

recrystallised cordierite. Phases were characterised using the SEM EDS det

free cotectic garnet (Grt1) is mantled by a retrograde growth rim (Grt2) s

breaks down to biotite2, sillimanite2/kyanite and garnet2 (reaction (5b)). In

orthopyroxene zone, PPL. (j) Static replacement of peak-metamorphic g

cooling. Cordierite is largely pinitised. Sample MM53, orthopyroxene zone

and cordierite breakdown. Local formation of sillimanite and biotite2 sugge

PPL. (l) Static retrograde replacement of sillimanite1 by muscovite. Biotite1supply according to reaction (5e). Sample 526/2; 250 m west of Loc. 52683 (2005) 143181Chessboard patterns occur along the western margin of

the biotitemuscovite zone and in all higher-grade

zones of the Sierra de Quilmes. The western margin of

the biotitemuscovite zone is also defined by musco-

vite breakdown reactions (see below).

Grt

Pin

Crd

Sil+Bt2

1mm

Ms

Sil Ms

Ms

Bt1

Bt1

1 mm

one. The mineral assemblage is QtzChlIllopaque mineralFKfs.on intersecting the horizontal Puncoviscana compositional layering

ngated quartz grains indicates absence of major plastic deformation.

he eastern garnetcordieritesillimanite zone. Sample 510, PPL. (d)

eosome (QtzN) is slightly deformed as indicated by formation of

. Random orientation of biotite in the neosome (BtN) indicates static

finer-grained quartz, biotite, and plagioclase. Sample 528A, cross-

the orthopyroxene zone. Biotite1 is part of the melano- or mesosome.

at the lower right shows irregular shape due to retrograde solid state

(Fig. 7). Note euhedral magmatic plagioclase inclusions in garnet.

ovite (MsC) and poikiloblastic cordierite (Crd) overgrowing the low-

itisation of cordierite is associated with hydrothermal activity in the

matic replacement of cordierite by sillimanite and muscovite in a

rade sillimanite, staurolite, magnetite and quartz along boundaries of

ector. Garnetcordierite zone, SEM image, sample 496. (i) Inclusion-

howing inclusions of sillimanite and biotite. Cordierite and biotite1a later stage, cordierite breaks down to pinite (Pin). Sample MM26B,

arnet and cordierite by sillimanite and biotite2 during near-isobaric

, PPL. (k) Retrograde formation of staurolite and kyanite by biotite1sts a retrograde PT path crossing the SilKy isograd. Sample 532A,

does not show evidence for breakdown, suggesting fluid-assisted K+

, CPL.

ithosDespite the large temperature range of ~440650

8C within the biotitemuscovite zone the rocks areuniform in their mineral assemblage. The mineral

assemblage and microstructures clearly indicate upper

greenschist and amphibolite facies temperatures of

metamorphism. The relatively low Al2O3 content of

~15 wt.% in the bulk composition of the Puncoviscana

sediments (Becchio et al., 1999; Lucassen et al., 2001),

which are greywackes rather than pelites (Taylor and

McLennan, 1985), prevented the formation of pyro-

phyllite. Consequently, neither chloritoid nor staurolite

or garnet was formed during prograde greenschist and

amphibolite facies metamorphism (Spear, 1993).

A PT estimation for rocks along the boundary of

the biotitemuscovite and the garnetcordieritesilli-

manite zone (Fig. 2) is based on mineral reactions and

field and thin section observation. In the field, the

muscovite-out reaction

Ms Qtz Kfs Sil1 V; 2the appearance of leucosomes (Fig. 4b) according to

the reaction

Ms=Kfs Pl Qtz V L; 3aand the low-/high-quartz transition (on the evidence of

chessboard subgrain patterns; Kruhl, 1996) were

mapped in an approximately 400 m wide zone. The

isograds defining the three mineral reactions are

closest to each other in PT space at 360 MPa and

650 8C (Fig. 5a). We assume that these PTconditions, with an arbitrarily chosen uncertainty of

F30 8C and F50 MPa, are the conditions of themetamorphic peak along the boundary of the biotite

muscovite and the garnetcordieritesillimanite zone.

3.3. The garnetcordieritesillimanite zone

The garnetcordieritesillimanite zone is charac-

terised by the formation of high-temperature mineral

assemblages, anatexis and large-scale modification

and overprinting of the PVCL. Increasing maximum

temperature of metamorphism in the southern and

western part of the study area led to increasing partial

melting. The characteristic feature of these migmatites

is the separation of mesosome and neosome, forming

a usually parallel but locally anastomosing migmatitic

S.H. Buttner et al. / Llayering (Fig. 4bd). Biotite-rich melanosomes are

rare but occur locally along the eastern zone margin.Rafts and layers of non-anatectic plagioclasebiotite

gneisses (Fig. 4c) of up to several meters in size are

present in both the garnetcordieritesillimanite and

the orthopyroxene zones. The gneissic rafts show the

same mineral assemblage, texture, and layering as do

the amphibolite facies rocks in the western part of the

biotitemuscovite zone, indicating that all three zones

derive from identical protoliths, i.e., the metasedi-

ments of the Puncoviscana Formation. In both, the

garnetcordieritesillimanite and the orthopyroxene

zones, intercalated Kfsporphyric granites are com-

mon. Their boundaries to the migmatites are diffuse,

irregular, and gradational. The granite bodies vary in

size from several meters to several hundred meters.

Within the garnetcordieritesillimanite zone

increasing grade of metamorphism from northeast to

southwest is documented by (1) the formation of

sillimanite, (2) the formation of QtzPlFKfs-bearingleucosome, (3) cordierite and/or garnet in leucosome,

and (4) progressively increasing grain size in the

leucosome reaching 11 cm (cordierite) to 18 cm

(garnet) in diameter.

Most likely, the muscovite breakdown and silli-

manite forming reaction at the northeastern margin of

the zone is reaction (2) (see above), forming the new

mineral assemblage sillimanite1, biotite1, K-feldspar,

plagioclase, and quartz (Fig. 3c). No relict textures of

prograde white mica breakdown have been observed

in thin section, suggesting that the reaction went to

completion. The paragenesis sillimanite1biotite1 is

common in the less anatectic eastern part, but occurs

locally throughout the zone, possibly as a metastable

relict. Occasionally, cordierite overgrows biotite and

sillimanite, suggesting the reaction

Bt1 Sil1 Qtz Crd Kfs V: 3bAnatexis is indicated by the presence of leucosome

veins up to 5 mm thick and locally folded (Fig. 4b),

and melt pockets of mm to cm in size, which become

more abundant and larger going westward away from

the eastern zone margin. The easternmost veins are

garnet and cordierite-free and contain quartz, plagio-

clase, and some K-feldspar suggesting reaction (3a).

The magmatic origin of the veins is indicated by the

euhedral shape of the feldspars (Vernon, 1986).

Magmatic textures in folded veins suggest syn-

83 (2005) 143181 151magmatic folding. Veins and leucosome layers with

the assemblage QtzPlFKfs occur also towards the

west of the garnetcordieritesillimanite zone and in

the orthopyroxene zone, suggesting that reaction (3a)

took place in all migmatites during crustal heating but

became less important with increasing temperature or

after the consumption of muscovite.

In the leucosome veins and layers further towards

the west, cordierite and garnet become more common.

Depending on the melt forming reaction and the Fe/

Mg ratio in the protolith and in the magma either

garnet or cordierite, or both can be present. The latter

garnet, plagioclase, and quartz (Fig. 3e), and in most

cases with K-feldspar. The mesosome contains biotite,

plagioclase and quartz. Nowhere are biotite1, garnet1,

and quartz in contact, suggesting biotite breakdown

according to the reactions:

Bt1 Grt1 Qtz V Opx L 4a

Bt1 Qtz V Opx L 4b

hlorit

onspi

CL,

igmat

pres

zone.

stals

ovisc

ne. S

) abo

S.H. Buttner et al. / Lithos 83 (2005) 143181 153assemblage is more common in the central and

western part of the garnetcordieritesillimanite zone.

Occasionally, K-feldspar is absent as a cotectic phase,

suggesting fluid-present melting. We address the

significance and relative importance of fluid-present

or fluid-absent melting in the migmatites of the Sierra

de Quilmes metamorphic complex at the end of

Section 3.4. Cordierite and/or garnet suggests break-

down of sillimanite and biotite by the reactions:

Bt1 Sil1 Qtz Grt1 Crd Kfs L 3cBt Sil Qtz Pl V Grt1 L 3dBt Sil Qtz Pl Grt1 Kfs L 3eBt1 Sil1 Qtz V Grt1 L 3f

Most of these reactions indicate minimum temper-

atures of metamorphism in excess of 650750 8C.Fig. 5a shows the positions of the isograds, the

literature sources are referenced in the figure caption.

3.4. The orthopyroxene zone

The westernmost zone in the metamorphic com-

plex of the Sierra de Quilmes contains orthopyroxene

in leucosomes, usually coexisting with cordierite,

Fig. 4. (a) Low-grade metamorphic Puncoviscana Formation in the c

interbedding of greywacke and psammo-pelitic layers (PVCL) less c

Antonio valley (see Fig. 2). (b) Magmatic veins crosscutting the PV

garnetcordieritesillimanite zone. 250 m northeast Loc. 472. (c) M

escaped along syn-magmatic shear zones (msz). Relics of PVCL are

the garnetcordieritesillimanite zone, close to the orthopyroxene

Leucosome layers contain mainly garnet. Occasional cordierite cry

metamorphic cordierite in psammo-pelitic/pelitic layers of the Punc

Cafayate, eastern margin of the Cafayate pluton. (f) Ductile shear zo

top-west) displacement. Loc. 540. (g) Two ductile shear zones (SZOutside the shear zones the migmatitic fabrics remain largely undefo

plagioclase and quartz show only minor anatexis or none.Bt1 Qtz Pl Opx Kfs L 4c

Bt1 Grt1 Qtz Opx Crd Kfs L 4dIn greywackes and pelitic systems (XMg close to

0.5) and in the stability field of sillimanite the

reactions (4a) and (4b) take place between ~780 8Cand ~820 8C (Vielzeuf and Holloway, 1988). At mid-crustal level, the fluid-absent reactions (4c) and (4d)

indicate minimum temperatures of ~750 8C (Clemensand Wall, 1981), and ~830 8C (Spear et al., 1999),respectively.

Apart from the mineral assemblage, the migmatites

of the orthopyroxene zone are similar to those in the

garnetcordieritesillimanite zone, although the

degree of partial melting and the proportion of

unfoliated layers of syn-migmatitic granite is generally

higher. Locally, these granites contain orthopyroxene.

In this area cordierite is less common or absent, which

possibly marks pressure-related cordierite breakdown.

The peak-metamorphic mineral assemblages in the

migmatites of the Sierra de Quilmes can be explained

by either fluid-present (e.g., reactions (3a), (3d), and

(4a)) or fluid-absent partial melting (e.g., reactions

(3c), (3e), and (4c)). Several studies (e.g., Stevens and

Clemens, 1993; Clemens and Droop, 1998) have

e zone. The absence of metamorphic biotite makes the sedimentary

cuous than in zones of higher-grade metamorphism. Loc. 551, San

and gathering of magma in melt pockets, are characteristics of the

ite from the central garnetcordieritesillimanite zone. Magma has

erved in rafts. Upper San Antonio Valley. (d) Layered migmatite in

A calcsilicate lens is oriented parallel to the migmatitic layering.

reach up to 4 cm diameter. 250 m east of Loc. 557. (e) Contact

ana Formation. Biotitemuscovite zone, 5 km southsouthwest of

C-fabrics, CV shear bands, and j-clasts indicate top-to-the right (i.e.ut 1 cm thick overprint the migmatitic layering in sample MM61.rmed. Small mesosome (Mes) domains of fine-grained biotite,

shown that fluid-excess in high-grade metamorphic

terrains is unlikely and that hydrous fluids, if present,

are immediately dissolved in the melt, lowering aH2O.

Therefore, fluid-present melting reactions are assumed

to be less efficient than fluid-absent melting. However,

in the Sierra de Quilmes the production of excess

hydrous fluids is obvious by the prograde breakdown

of chlorite and muscovite. These reactions can be

inferred to have taken place in the migmatitic zones

during crustal heating, because the migmatites derive

from the same protoliths as the biotitemuscovite zone

and the chlorite zone, where solid state dehydration

reactions are obvious. Partial solid-state biotite break-

does not crystallise from granitic/granodioritic melt at

temperatures above ~720 8C (Whitney, 1975, 1988).This temperature was exceeded in most migmatites of

the Sierra de Quilmes. We assume that melt produced

by fluid-present melting reactions was removed from

the source at, or shortly after, the thermal peak, and

crystallised as intercalated Kfsporphyric granites.

The diffuse and gradational contacts of the intercalated

granites indicate the close relationship of granite

formation and anatexis. At a larger scale, the formation

of the Cafayate granite can be seen in this context (see

below). After consumption of the excess hydrous fluid,

anatexis continued by fluid-absent melting (reactions

memb

~475

atite

nterpr

etrog

aximu

onstra

2A: g

haded

int oc

4b), (

eege (

tiphas

S.H. Buttner et al. / Lithos 83 (2005) 143181154down (3b) might have supplied additional fluid during

crustal heating or at the thermal peak in the high-grade

metamorphic zones. It seems to be likely that the fluid

phase remained close to its source and was available

for fluid-present partial melting reactions at least in the

early stages of anatexis. Two observations support this

assumption: (1) Evidence for the efficient removal of

fluids, such as abundant hydrothermal veins in the

non- or weakly anatectic parts of the metamorphic

complex, is absent. (2) K-feldspar, or another potas-

sium-bearing mineral, is frequently absent in cordier-

ite-, garnet-, or orthopyroxene-bearing leucosomes,

although the melting reaction was biotite breakdown.

This shows that K-feldspar was not always a cotectic

phase, which fluid-absent melting reactions would

have produced (e.g., (3c), (3e), and (4d)). It is obvious

that potassium-rich melt has been removed after

crystallisation of the cotectic mineral (Crd, Grt,

Opx), and quartz and plagioclase. At mid-crustal

levels and moderate or high water content, K-feldspar

Fig. 5. (a) PTt evolution of medium- and high-grade metamorphic

isobaric cooling of the migmatites after peak metamorphism at

deformation ceased at amphibolite facies temperature before the migm

data calculated from core compositions of the peak assemblage, i

calculated from rim compositions of peak metamorphic minerals or r

of mid-crustal migmatites in the Sierra de Quilmes. We assume m

However, most calculated temperatures are in agreement with and c

Samples 557G and MM61: orthopyroxene zone. Samples 506 and 53

zone in the orthopyroxene zone. 513A: biotitemuscovite zone. S

between biotitemuscovite and garnetcordieritesillimanite zone (jo

(2): Spear and Cheney (1989); (3a): Holland (1979); (3b), (3f), (4a), (

(1999); (3d), (3e), and (4c): Clemens (1984); (8): van Groos and H

Massonne and Schreyer (1987). (b) (f) PT calculations using mulSolid lines show equilibria of the peak assemblage, dashed lines those

conventional geothermobarometry are indicated in (c), (d), and (e).(3c), (3e), (4c), and (4d)), producing cotectic K-

feldspar, which is present in many leucosomes. The

importance of fluid-absent or fluid-present partial

melting has varied locally and temporally in depend-

ence of fluid availability, pressure, and temperature.

3.5. The Cafayate pluton and its contact

metamorphism

The composite Cafayate pluton intruded roughly

along the boundary between the biotitemuscovite and

the garnetcordieritesillimanite zone, hence, at the

transition between non-anatectic and anatectic Punco-

viscana metasediments (Fig. 2). The pluton consists of

at least four different granitoids ranging in composi-

tion from granites to tonalites. The different types

appear to be peraluminous (cordierite- and sillimanite-

bearing) to metaluminous (epidote-bearing). More

detailed petrographic and geochemical descriptions

have been published previously (Rapela, 1976b;

ers of the Sierra de Quilmes metamorphic complex. Retrograde near-

Ma is accompanied by non-penetrative ductile shearing. Ductile

s reached the stability field of kyanite at ~440 Ma (see text). C : PT

eted as peak metamorphic conditions; R : retrograde equilibration

rade assemblages. The arrow indicates the approximate cooling path

m uncertainties of F100 MPa and F50 8C for PT calculations.ined by critical mineral assemblages and petrogenetic grids as well.

arnetcordieritesillimanite zone. 542: Mylonite from a ductile shear

circular field: metamorphic peak conditions along the boundary

currence of reactions (2), (3a), and (8)). (1): Simpson et al. (2000);

6) and (7): Vielzeuf and Holloway (1988); (3c) and (4d): Spear et al.

1973); Al2SiO5 phase diagram: Holdaway (1971); Si in muscovite:

e equilibria (TWEEQU 1.02 and 2.02; Berman, 1988, 1991, 1992).of the retrograde stage (details are given in the text). Results of

KyAnd

Chl K

fsBt

Ms

Qtz

V[1]

Si (Ms)=3

.12

Si (Ms

)=3.08

reg. met.513 A

cont. met.

KySil

low

-Qt

zhi

gh-Qt

z

SilAnd

400 500 600 700 800 900T [C]

[2]

[3f]

200

400

600

800

P [M

Pa]

[6]

[8]

~475 Ma

[4d]

[4b]

[7][3c

]

[2] Ms+Qtz=Kfs+Sil+V[3a] Ms/Kfs+Pl+Qtz+V=L [3b] Bt+Sil+Qtz=Crd+Kfs+V[3c] Bt+Sil+Qtz=Grt+Crd+Kfs+L[3d] Bt+Sil+Qtz+Pl+V=Grt+L[3e] Bt+Sil+Qtz+Pl=Grt+Kfs+L[3f ] Bt+Sil+Qtz+V=Grt+L[4a] Bt+Grt+Qtz+V=Opx+L[4b] Bt+Qtz+V=Opx+L[4c] Bt+Qtz+Pl=Opx+Kfs+L[4d] Bt+Grt+Qtz=Opx+Crd+Kfs+L [6] Bt+Crd+Qtz+V=Grt+L [7] Bt+Crd+Qtz+V=Opx+L [3d]

[2]

[3a ]

[3a]

[4a]

[3e]

[4c]

532AR

??

557GC

532AC

542R

506C

542C

~475 Ma

~475 Ma

~440 Ma

506R

MM61C

MM61R ??

557GR

[3b]

100

300

500

700

300 500 700 900

900

P [M

Pa]

T [C]

Sample 506

6An+2Pyp

+3Qtz

2Gross+3C

rd

Gros

s+Qt

z+2P

yp3C

rd

Alm

+Ph

lPy

p+An

n

6Sil+

5Gro

ss+3C

rd+

2Ann

2Alm

+15

An+2P

hl

[1] [2]

[1] 2Alm+2Phl+5Qtz+4Sil=3Crd+2Ann[2] 6Sil+5Gross+3Crd+2Ann=2Alm+15An+2Phl[3] 2Alm+4Ky+2Phl+5Qtz=3Crd+2Ann

4Sil+5Qtz+2Pyp3Crd

Alm

+Ph

lPy

p+An

n

4Ky+2Py+

5Qtz

3Crd

[3]

a

b


100

300

500

700

300 500 700 900

4Sil+5Qtz+2Pyp

3Crd3Q

tz+2Pyp+6A

n

3Crd

+2Gross

6Sil+

5Gro

ss+3C

rd

15An

+2P

yp

Gros

s+Qt

z+2S

il

3An

900P

[MPa

]

66Sil+10Pyp+8Alm

+12V

15Crd

+6St

8Alm+46Ky+12V

6St+25Qtz

12St+

65Qtz

+46P

yp

16Alm

+69C

rd+24

V

5Qtz+4K

y+2Pyp

3Crd

5Qtz+

2Pyp+

4Ky

3Crd

6St+

25Qt

z8A

lm+

69Cr

d+24

V

Sample 532A

[1]

[1] Alm+Phl=Pyp+Ann (Holdaway et al. 1997)

9Qtz+6Pyp+4Ann

3Crd+6Fsl+4Phl

Alm

+Ph

lPy

p+An

n

6Alm

+2Ph

l+9Qt

z

6Fsl+

3Crd+

2Ann

4Alm+

2Pyp+9

Qtz

6Fsl+3C

rdSample 557G

T [C]

300 500 700 900 T [C]

Alm

+Ph

lPy

p +An

n

6An

+2Pyp+3Q

tz

2Gross+3C

rd

[1] 2Qtz+2Phl+6An+2Alm=2Ann+3Crd+2Gross[2] Alm+Phl=Pyp+Ann

(Holdaway et al., 1997; Holdaway, 2000)[3] Crd-Bt thermobaromerty

(Holdaway and Lee, 1977)

[1]

[2]

[3]

c

100

300

500

700

900

P [M

Pa]

d

Fig. 5 (continued).


ithose 900

S.H. Buttner et al. / LRapela et al., 1998b, 1990). Parts of the contacts

between different granitoids are irregular and diffuse,

indicating their coexistence as magmas; others are

300 500

300 500

Sample 542

[1]

[1] Gross+Qtz+2Sil=2An[2] Alm+Phl=Pyp+Ann

100

300

500

700

P [M

Pa]

100

300

500

700

P [M

Pa]

f

Sample MM61

H: Bt-Grt thermometry and GASP barometry (Holdaway et al., 1997; Hol

Fig. 5 (contiross

83 (2005) 143181 157straight, suggesting intrusion into solid pre-existing

parts of the pluton. At the eastern, northern, and

southern pluton margins, the intruding magma cross-

700 900

700 900 T [oC]

T [oC]

3Sil+K

fs+2G

ross

+Alm+

V

Ann+

6An

Sil+Q

tz+Gr

oss

3An

Kfs+

Alm

+V

Ann+

Qtz+

Sil

2Alm+G

ross+Kfs+2V

3Qtz+2An+2Ann

[2]

[2]

2Sil+

Qtz+

G3A

n

Phl+

Alm

Ann+

Pyp

Phl+

Alm

Ann

+Py

p

2Sil+

Qtz+

Gros

s3A

n

H

H

daway, 2000)

nued).

Lithoscuts the Puncoviscana metasediments litparlit with

sharply defined and discordant contacts. Particularly

along the southern margin the intrusive magma cuts

the migmatitic layering. In contrast, the western

contact is partly diffuse with gradual transitions from

intrusive granite to in situ migmatite, suggesting the

coexistence of plutonic and anatectic melt.

Macroscopically conspicuous contact metamor-

phism formed hornfels, cordierite and tourmaline

schists at the eastern, northern, and locally at the

southern margin of the pluton. Tourmaline (13 mm

length), cordierite (13 cm diameter; Fig. 4e),

muscovite and biotite (12 mm) are the most common

contact-metamorphic minerals. There is no such

evidence for contact metamorphism along the western

contact, where the pluton grades into layered migma-

tites, indicating near-peak PT conditions of the

country rock during pluton emplacement.

Contact-metamorphic minerals overgrow the pre-

existing layering and the low-strain foliation of the

Puncoviscana metasediments of the biotitemuscovite

zone (Figs. 3f and 4e). Locally, regional metamorphic

biotitemuscovite assemblages show grain coarsening

or are overgrown by static new micas with slightly

different composition (see Section 4). These new

micas are overgrown by large cordierite crystals,

possibly reflecting polyphase intrusions leading to

several generations of contact minerals.

3.6. Pegmatites

Pegmatites occur as three distinct types. The oldest

group of pegmatites has been seen only in the

orthopyroxene zone and consists of disrupted, centi-

metre- to meter-thick leucocratic dykes and lenses

with diffuse margins, frequently intruding parallel to,

but locally also discordant to the migmatitic layering

of the country rock. This suggests a formation of the

pegmatitic melt during regional anatexis and emplace-

ment during or shortly after the formation of the

migmatitic layering. These pegmatites consist of

quartz, K-feldspar, and plagioclase. Tourmaline is

rare and white mica is absent.

A second, younger group of pegmatites (type 2)

crosscuts its country rock along sharp contacts. They

are most common in the garnetcordieritesillimanite

S.H. Buttner et al. /158and biotitemuscovite zones but occur in all metamor-

phic zones of the complex except the northern part ofthe chlorite zone. This group of pegmatites occurs as

near-vertical dykes and dyke swarms. Their thickness

varies between several centimetres and several meters.

Most of them strike NNWSSE, parallel to the long

axis of the pluton and the regional metamorphic

layering. A minority strikes EW. The pegmatites are

coarse- to very coarse-grained and contain K-feldspar,

plagioclase, quartz, muscovite,Ftourmaline,Fapatite,Fgarnet, andFzircon. Some show evidence of plasticdeformation but most are undeformed and cut ductile

shear zones (see Section 3.7). Magmatic quartz of

isometric or irregular shape is weakly deformed with

coarsely sutured grain boundaries, suggesting minor

grain boundary migration but no recrystallisation (i.e.,

no formation of new grains). Chessboard patterns, as

well as prism-parallel subgrain boundaries are present,

suggesting that the pegmatites crystallised at temper-

atures near the high-/low-quartz transition (Kruhl,

1996). Their crosscutting relationship with the mig-

matitic layering and the shear zones indicates that the

pegmatites postdate the anatexis and most of the

ductile deformation in their host rocks. However, the

formation of the pegmatitic melt might be related to

continuing anatexis or plutonism at crustal levels

deeper than the orthopyroxene zone. The persistence

of high temperature until the upper Silurian (~ 420

Ma) in the basement of the northern Sierras Pampea-

nas has been shown by Lucassen and Becchio (2003).

A third group of pegmatites (type 3) intruded into

the Cafayate pluton. These pegmatites are petrograph-

ically similar to the second group, but show transitional

features to miaroles, suggesting a genetic relationship

with the Cafayate pluton. We have dated type 2

(samples 472B, 552, MM72, and MM110) and type 3

pegmatites (samples 515B and 526; see Section 5).

3.7. Shear zones and retrograde metamorphism

The migmatites of the garnetcordieritesillimanite

and orthopyroxene zones are affected by retrograde

mineral reactions, which are partly associated with

syn- to post-migmatitic ductile shear zones. Lower-

grade metamorphic zones of the complex are much less

deformed. In the migmatites, the earliest increment of

deformation is coeval with the production of melt, as

indicated by the escape of magma along syn-magmatic

83 (2005) 143181shear zones in meta- and diatexites (Fig. 4c) and by

syn-magmatic folding of leucosome veins (Fig. 4b).

Hence, the onset of deformation occurred at the

thermal peak. Deformation in the migmatites contin-

ued under high-grade metamorphism but in the solid

state. Non-penetrative east-dipping ductile shear zones

are oriented parallel to the migmatitic layering and

show top-to-the west or top-to-the northwest sense of

shear (Fig. 4f). The shear zone thickness ranges from a

few centimetres to ~300 m. Ductile shearing also

affected the Cafayate pluton. The shear zones overprint

the migmatitic or magmatic mineral assemblage by

recrystallisation and pressure solution, the latter

indicated by enrichment of biotite in cleavage

domains. Feldspar and cordierite show recrystallisa-

tion. In other shear zones, cordierite is replaced by

sillimanite (Fig. 3g). Recrystallised cordierite shows

fine-grained staurolite, sillimanite2, quartz, and mag-

netite along the grain boundaries (Fig. 3h), indicating

amphibolite facies temperatures of deformation.

Due to the non-penetrative style of deformation

most migmatites do not show significant plastic

deformation, and retrograde mineral assemblages with

static textures and random mineral orientation are

typical in these rocks. The retrograde metamorphic

overprint is strongest in cordierite and garnet

cordierite-bearing assemblages. Common reactions

are inferred from replacement textures:

Opx Kfs V Bt2 QtzFMag 5a

Crd Bt1 Grt2 Bt2 Sil2=KyFV 5b(Fig. 3i)

Grt1 Crd V Bt2 Sil2 5c(Fig. 3j)

Bt1 Crd St Bt2 Sil2=KyFV 5d(Fig. 3k)

Sil1 V Ms 5e(Fig. 3l).

Garnet2, formed by reaction (5b), does not form

individual crystals but poikilitic rim zones around

garnet1, with small biotite, sillimanite, or quartz

inclusions (Fig. 3h). The greenish biotite2 has slightly

A

2

il/Ky

d bio

506B

S.H. Buttner et al. / Lithos 83 (2005) 143181 159F

Bt12.0-3.5 wt% TiO

St0.5-1.4 wt% ZnO

S

Grt

Bt1+Crd=St+Sil/Ky+Bt2

532A

Fig. 6. AFM diagram for the retrograde breakdown of cordierite an

cooling from granulite to amphibolite facies temperatures. Sampleshatched and shaded fields show the retrograde assemblage of samples 532

contribution of garnet to the retrograde formation of staurolite and biotiteM

+ Qtz+ Kfs+ H2O

Bt20.0-1.1 wt% TiO2

black symbols: sample 506white symbols: sample 532A

Crd

506

tite1 to biotite2, aluminosilicate and staurolite during near-isobaric

and 532A come from the garnetcordieritesillimanite zone. TheA and 506, respectively. Petrographic evidence does not suggest the

2 (Fig. 3k).

higher XMg than biotite1 and significantly lower TiO2content (Fig. 6). Garnet1 in the crystal core is

compositionally identical to euhedral cotectic garnet1from other samples (Fig. 7). Garnet2 of the rim zone

shows an increase in almandine and decrease in

pyrope contents compared with garnet1. As sillimanite

is absent at peak conditions, and the core of the garnet

had formed during the magmatic stage, it is unlikely

that the rim-proximal sillimanite inclusions are

inherited relicts of the prograde stage. Thus, we

interpret this texture as retrograde growth of garnet2 at

the expense of biotite1 and cordierite, according to

reaction (5b). Occasionally, spessartine contents

increase slightly towards the rim (sample 532A; Fig.

7). Retrograde garnet resorption cannot be excluded in

these cases. However, manganese fractionation during

retrograde garnet growth might be equally possible as

manganese becomes available from cordierite break-

down (reaction (5b)). Cordierite contains up to 0.5

wt.% MnO (Table 1 and Electronic Data Supplement).

All cordierite breakdown reactions produce amphib-

olite facies index minerals, similar to the reactions

observed in the retrograde shear zones. Kyanite occurs

exclusively in static environments or post-kinemati-

cally in shear zones, suggesting that deformation had

ceased when the PT conditions of the migmatites

reached the stability field of kyanite.

4. Mineral chemistry and geothermobarometry

Microchemical analyses were obtained using a

wavelength dispersive CAMECA Cambax Microbeam

(PAP/XMAS correction) at the Technische Universit7tBerlin and a JEOL JXA 733 Superprobe (ZAF

correction) at Rhodes University. We used an accel-

erating potential of 20 kVand a beam current of 1518

nA. Beam widths were 5 Am for garnet and orthopyr-oxene analyses and 10 Am for feldspar, cordierite, andmica analyses. Ferric iron has been calculated for

garnet and orthopyroxene according to charge balance

and lattice site occupation. Fe3+ concentrations are

core rimCrd

350 m557G

25

MM 61

40

Alm -

Prp

Sps

Grs

And

XMg*

rim rim 1.6 mm core

PlQtz


5

10

15

20

25

30

35

40

1 3 5 7 9 11 13 15 17 19 21 23

rim rim 5 mm core

PlBt 532A

mol%

mol%

0

5

10

15

20

25

30

35

1 4 876532 9 10 11 12

Grt2Grt1

Grt2

Grt2Grt1Fig. 7. Zonation patterns of garnet from the high-grade metamorphic zones

and interpretation see Section 4.1 3 5 7 9 11 13 15 17 190

5

10

15

20

mol%0

5

10

15

20

25

30

35

40

30

35

40rim rim 900 m core

Bt Bt542

1 2 3 4 5 6 7

40

100

Grt1

Grt2Grt2

Grt1

mol%. Mineral phases in contact with garnet are indicated. For discussion

Table 1

Representative electron microprobe analyses

Sample Biotite

MM61 MM61 MM61 557G 557G 557G 532A 532A 506A 506A 513 513

Bt1 Bt2 Bt2 R Grt C I Opx R Crd Bt1 Bt2 Bt1 Bt2 rm cm

SiO2 35.83 35.86 35.89 34.87 34.68 35.51 35.26 37.20 36.25 38.00 36.06 35.77

TiO2 4.84 4.87 5.24 5.49 5.34 4.83 2.40 0.58 2.56 0.04 1.47 1.79

Al2O3 16.75 17.08 16.89 15.79 15.47 16.52 18.09 19.01 16.32 18.47 18.93 18.80

MgO 10.51 10.83 12.69 10.29 10.39 11.65 11.98 12.35 11.62 14.03 10.82 10.60

CaO 0.00 0.00 0.00 0.08 0.03 0.02 0.00 0.09 0.03 0.03 0.02 0.05

MnO 0.01 0.08 0.02 0.13 0.10 0.00 0.03 0.15 0.30 0.17 0.12 0.14

FeO 17.84 17.09 14.39 18.01 17.94 16.30 16.51 15.37 19.36 14.36 17.48 17.89

Na2O 0.01 0.04 0.03 0.07 0.04 0.07 0.29 0.31 0.29 0.46 0.33 0.32

K2O 8.32 8.41 8.48 9.56 9.25 9.36 9.63 8.90 9.66 9.42 8.82 9.03

F 0.58 0.67 0.77 1.01 0.90 1.28 0.49 0.54 0.33 0.63 0.40 0.37

Total 94.75 94.93 94.52 95.29 94.13 95.53 94.68 94.58 96.71 95.59 94.46 94.76

Cations: O=22

Si 5.39 5.36 5.32 5.38 5.40 5.40 5.38 5.59 5.48 5.64 5.49 5.45

Ti 0.55 0.55 0.58 0.64 0.63 0.55 0.28 0.07 0.29 0.00 0.17 0.20

Al 2.97 3.01 2.95 2.87 2.84 2.96 3.25 3.37 2.91 3.23 3.39 3.37

Mg 2.36 2.41 2.80 2.36 2.41 2.64 2.72 2.77 2.62 3.10 2.45 2.41

Ca 0.00 0.00 0.00 0.01 0.01 0.00 0.00 0.02 0.00 0.00 0.00 0.01

Mn 0.00 0.01 0.00 0.02 0.01 0.00 0.00 0.02 0.04 0.02 0.02 0.02

Fe 2.24 2.14 1.78 2.32 2.34 2.07 2.11 1.93 2.45 1.78 2.22 2.28

Na 0.00 0.01 0.01 0.02 0.01 0.02 0.09 0.09 0.08 0.13 0.10 0.09

K 1.59 1.60 1.60 1.88 1.84 1.82 1.87 1.71 1.86 1.78 1.71 1.75

F 0.27 0.32 0.36 0.49 0.44 0.62 0.23 0.25 0.16 0.30 0.19 0.18

Total 15.38 15.40 15.43 15.99 15.92 16.10 15.93 15.81 15.74 15.70 15.75 15.76

XMg 0.51 0.53 0.61 0.50 0.49 0.44 0.56 0.59 0.52 0.64 0.52 0.51

Sample Garnet

MM61 MM61 557G 557G 532A 532A 542 542 506A 506A

R C C R Crd C R Bt C R Bt Grt1 Grt2

SiO2 38.17 38.48 37.20 37.09 37.38 37.57 38.52 37.39 37.26 37.25

TiO2 0.00 0.02 0.00 0.00 0.00 0.00 0.00 0.14 0.03 0.00

Al2O3 21.79 21.94 21.24 21.47 21.60 21.28 21.61 21.93 21.39 21.34

Cr2O3 0.01 0.00 0.02 0.07 0.00 0.00 0.02 0.05 0.00 0.00

Fe2O3 0.00 0.00 1.00 0.60 0.00 0.00 0.00 0.00 0.80 0.00

MgO 6.25 7.35 6.36 5.84 5.66 3.53 6.05 4.07 4.89 3.65

CaO 1.23 1.23 1.05 1.01 1.02 1.06 1.62 1.49 1.46 0.99

MnO 1.84 1.88 2.11 1.34 2.48 3.74 1.58 1.39 4.02 3.38

FeO 31.23 29.93 30.54 31.94 30.88 32.91 31.25 32.97 30.51 33.18

Total 100.51 100.83 99.52 99.37 99.02 100.08 100.66 99.43 100.35 99.80

Cations: O=24

Si 5.98 5.97 5.91 5.91 5.96 6.02 6.02 5.97 5.92 5.99

Ti 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00

Al 4.02 4.01 3.98 4.03 4.06 4.02 3.98 4.13 4.00 4.04

Cr 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.01 0.00 0.00

Fe3+ 0.00 0.00 0.12 0.07 0.00 0.00 0.00 0.00 0.10 0.00

Mg 1.46 1.70 1.51 1.39 1.35 0.84 1.41 0.97 1.16 0.88

Ca 0.21 0.20 0.18 0.17 0.17 0.18 0.27 0.25 0.25 0.17

(continued on next page)


Sample Garnet

MM61 MM61 557G 557G 532A 532A 542 542 506A 506A

R C C R Crd C R Bt C R Bt Grt1 Grt2

Mn 0.24 0.25 0.28 0.18 0.34 0.51 0.21 0.19 0.54 0.46

Fe2+ 4.09 3.88 4.06 4.26 4.12 4.41 4.09 4.40 4.05 4.46

Total 16.01 16.02 16.04 16.03 16.00 15.97 15.98 15.94 16.03 15.99

XMg 0.26 0.30 0.27 0.25 0.25 0.16 0.26 0.18 0.22 0.16

Sample Cordierite

557G 557G 532A2 532A2 506B 513

C R Grt C R Grt Rn cm

SiO2 48.14 48.91 49.04 48.38 49.29 48.39

TiO2 0.01 0.00 0.00 0.03 0.00 0.01

Al2O3 33.47 33.81 32.45 32.34 32.16 32.65

MgO 9.79 10.91 8.39 8.80 8.64 8.54

CaO 0.02 0.04 0.03 0.03 0.03 0.01

MnO 0.07 0.07 0.41 0.47 0.34 0.43

FeO 6.00 4.67 6.85 5.91 7.28 6.87

Na2O 0.05 0.10 0.27 0.17 0.26 0.55

K2O 0.02 0.05 0.00 0.01 0.00 0.01

Total 97.58 98.56 97.45 96.15 98.00 97.45

Cations: O=18

Si 4.95 4.95 5.06 5.04 5.07 5.01

Ti 0.00 0.00 0.00 0.00 0.00 0.00

Al 4.05 4.03 3.95 3.97 3.90 3.98

Mg 1.50 1.65 1.29 1.37 1.32 1.32

Ca 0.00 0.00 0.00 0.00 0.00 0.00

Mn 0.01 0.01 0.04 0.04 0.03 0.04

Fe 0.52 0.40 0.59 0.52 0.63 0.59

Na 0.01 0.02 0.05 0.03 0.05 0.11

K 0.00 0.01 0.00 0.00 0.00 0.00

Total 11.03 11.05 10.99 10.98 11.01 11.06

XFe 0.26 0.19 0.31 0.27 0.32 0.31

Sample Staurolite

532A 532A 506B

R C Rn

SiO2 36.96 39.73 26.75

TiO2 0.03 0.28 0.65

Al2O3 46.86 43.91 54.11

MgO 1.60 1.89 1.57

CaO 0.01 0.09 0.03

MnO 0.65 0.51 0.70

FeO 10.50 9.55 12.90

ZnO 0.67 0.52 0.84

Na2O 0.02 0.01 0.02

K2O 0.04 0.38 0.01

Total 97.34 96.87 97.57

Cations: O=23

Si 5.03 5.39 3.74

Ti 0.00 0.03 0.07

Al 7.51 7.02 8.92

Table 1 (continued)


Sample Staurolite

532A 532A 506B

R C Rn

Mg 0.32 0.38 0.33

Ca 0.00 0.00 0.00

Mn 0.07 0.06 0.08

Fe 1.19 1.08 1.51

Zn 0.07 0.05 0.09

Na 0.00 0.00 0.01

K 0.01 0.07 0.00

Total 14.22 14.10 14.73

XFe 0.79 0.74 0.82

Sample Plagioclase

C R

Crd

R

Opx

R

Grt

R

Opx/Grt

Rn

Opx

C C Rn R R C Rn R C C Rn R

557G 557G 557G 557G 557G 557G 557G 506A 506A 506A 532A1 532A1 532A1 542 542 MM61 MM61 MM61

SiO2 59.94 59.69 60.46 61.42 59.62 59.58 64.22 59.43 60.42 60.22 60.91 60.51 59.77 59.08 60.42 60.80 59.99 59.70

Al2O3 25.08 24.99 24.73 24.11 25.00 24.71 22.84 25.22 25.72 25.92 25.15 24.25 24.79 25.86 24.72 25.15 25.22 26.26

CaO 6.33 6.52 6.52 5.76 6.44 6.44 4.04 6.95 6.97 6.92 6.19 5.72 5.68 8.25 6.96 6.25 6.33 7.24

MnO 0.01 0.04 0.04 0.00 0.00 0.00 0.01 0.00 0.00 0.01 0.07 0.07 0.03 0.01 0.00 0.02 0.00 0.00

FeO 0.06 0.20 0.22 0.00 0.22 0.33 0.02 0.03 0.02 0.02 0.16 0.16 0.13 0.09 0.03 0.15 0.20 0.00

Na2O 7.58 7.48 7.41 8.64 7.34 7.49 0.00 7.67 7.68 7.16 7.94 8.22 8.25 7.02 7.67 7.46 7.57 6.82

K2O 0.26 0.25 0.29 0.05 0.26 0.26 9.42 0.13 0.14 0.22 0.05 0.05 0.05 0.16 0.13 0.05 0.09 0.07

Total 99.25 99.19 99.68 99.96 98.87 98.80 100.80 99.43 100.95 100.46 100.50 99.01 98.74 100.50 99.95 99.89 99.39 100.09

Cations: O=8

Si 2.69 2.68 2.70 2.73 2.68 2.69 2.82 2.66 2.67 2.66 2.69 2.72 2.69 2.63 2.69 2.70 2.68 2.65

Al 1.32 1.32 1.30 1.26 1.33 1.31 1.18 1.33 1.34 1.35 1.31 1.28 1.32 1.36 1.30 1.32 1.33 1.37

Ca 0.30 0.31 0.31 0.27 0.31 0.31 0.19 0.33 0.33 0.33 0.29 0.28 0.27 0.39 0.33 0.30 0.30 0.34

Mn 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Fe 0.00 0.01 0.01 0.00 0.01 0.01 0.00 0.00 0.00 0.00 0.01 0.01 0.00 0.00 0.00 0.01 0.01 0.00

Na 0.66 0.65 0.64 0.74 0.64 0.65 0.00 0.67 0.66 0.61 0.68 0.72 0.72 0.61 0.66 0.64 0.66 0.59

K 0.01 0.01 0.02 0.00 0.01 0.01 0.80 0.01 0.01 0.01 0.00 0.00 0.00 0.01 0.01 0.00 0.01 0.00

Total 4.99 4.99 4.98 5.01 4.98 4.99 5.00 5.01 5.00 4.97 4.99 5.00 5.01 5.00 4.99 4.96 4.98 4.96

An 31.1 32.0 32.1 26.8 32.2 31.7 19.1 33.4 33.4 34.8 30.0 27.7 27.5 39.0 33.2 31.7 31.6 36.9

Ab 67.4 66.5 66.2 72.9 66.3 66.8 79.8 66.1 66.1 64.4 69.6 72.0 72.2 60.1 66.1 68.1 68.0 62.8

Or 1.5 1.5 1.7 0.2 1.5 1.5 1.4 0.7 0.8 1.3 0.3 0.3 0.3 0.9 0.7 0.3 0.5 0.4

Sample Orthopyroxene

C R Bt1 R Grt Rn Grt

557G 557G 557G 557G

SiO2 46.99 46.78 49.17 48.59

TiO2 0.13 0.12 0.13 0.13

Al2O3 5.62 5.59 3.79 4.07

Cr2O3 0.06 0.08 0.02 0.00

Fe2O3 0.00 0.00 0.00 0.00

MgO 14.99 14.94 17.96 17.23

CaO 0.10 0.10 0.12 0.06

MnO 1.01 0.99 0.53 0.53

FeO 30.55 30.88 27.47 28.78

Na2O 0.03 0.04 0.01 0.00

Table 1 (continued)

(continued on next page)


R Bt1

557G

0.01

99.53

1.84

0.00

0.26

0.00

0.00

0.87

0.00

0.03

1.01

0.00

0.00

4.03

53.6

ntion

tioned

ct me

Lithosgenerally very low. In biotite the total iron is given as

FeO. Representative microprobe analyses are shown in

Sample Orthopyroxene

C

557G

K2O 0.00

Total 99.49

Cations: O=6

Si 1.84

Ti 0.00

Al 0.26

Cr 0.00

Fe3+ 0.00

Mg 0.88

Ca 0.00

Mn 0.03

Fe2+ 1.00

Na 0.00

K 0.00

Total 4.02

XFe 53.2

C: core composition; R: rim composition (neighbouring phase is me

rim, usually closer than 30 Am); I: inclusion (host phase can be menphase is mentioned); M: composition of a matrix crystal; cm: conta

Table 1 (continued)

S.H. Buttner et al. /164Table 1 and in the Electronic Data Supplement. The

essential petrographic characteristics of the samples

used for geothermobarometry are given below; more

information can be found in the Electronic Data

Supplement.

Temperatures and pressures were calculated using

analyses from minerals in contact with each other

using the multiphase equilibria (TWEEQU; Berman,

1991), and several methods of conventional geo-

thermobarometry. We have preferred the Berman

method where possible. For thermobarometry we

used both single analyses from co-existing mineral

grains and average compositions. The results obtained

from single analyses and average compositions show

only minor differences and, if not mentioned other-

wise, we present only the results obtained from

average mineral compositions. Average compositions

were calculated from at least five single analyses per

mineral phase involved. Unless otherwise noted we

have used the TWEEQU 2.02 and the BA95/BA96

(Berman, 1988; Berman and Aranovich, 1996), and

Fuhrman and Lindsley (1988) mixing models for

biotite, garnet, cordierite, orthopyroxene, and plagio-

clase, respectively. A water activity of aH2O=1 wasassumed and only stable reactions were calculated.

Petrologic equilibrium was assumed from textural

R Grt Rn Grt

557G 557G

0.02 0.02

99.21 99.42

1.90 1.88

0.00 0.00

0.17 0.19

0.00 0.00

0.00 0.00

1.03 1.00

0.00 0.00

0.02 0.02

0.89 0.93

0.00 0.00

0.00 0.00

4.01 4.02

46.1 48.3

ed, if applicable); Rn: brim-nearQ composition (analysis close to the); Pr: analysis in a profile (when rim composition, the neighbouring

tamorphic crystal; rm: regional metamorphic crystal.

83 (2005) 143181evidence, such as straight grain boundaries, foam, or

magmatic textures. More detailed petrographic and

chemical descriptions, detailed EPM analyses, and

GPS coordinates of sample locations are provided in

the Electronic Data Supplement.

In support of the PT estimations from mineral

assemblages and reactions, and petrogenetic grids

(Section 3) we have selected six samples from the

biotitemuscovite zone (513A), the garnetcordierite

sillimanite zone (506, 532A), and the orthopyroxene

zone (557G, MM61, and 542) for geothermobarom-

etry. The sample locations are shown in Fig. 2.

Sample 513A is a layered Puncoviscana metasedi-

ment of the regional metamorphic mineral assemblage

QtzMsBtPlKfs from the biotitemuscovite zone.

Contact metamorphic muscovite, biotite and cordierite

overgrew the foliation of the regional metamorphic

peak in the course of Cafayate pluton emplacement.

Using average mica compositions for biotitemusco-

vite geothermobarometry (Hoisch, 1989) in combina-

tion with phengite barometry (Massonne and

Schreyer, 1987), the regional metamorphic assem-

blage was determined to have equilibrated at 459

(F25) 8C and 280 (F100) MPa (Fig. 5a). The contact

ithosmetamorphic overprint occurred at slightly higher

temperature but similar pressure (532 8C (F25) and240 (F100) MPa).

Sample 506 comes from the central garnet

cordieritesillimanite zone west of Tolombon (Fig.

2). The leucosome consists of the assemblage QtzPl

KfsCrdFGrtFBt1. Biotite1 and sillimanite1 occur inthe mesosome. Biotites in the mesosome and in the

leucosome are compositionally identical. During the

retrograde PT evolution, cordierite- and biotite1-

breakdown formed staurolite and sillimanite2. Aver-

age core compositions of the leucosome minerals

CrdGrtPlQtzBt have been used to determine P

and T using multiphase equilibria, yielding the

intersection of seven independent mineral reactions

close to 540 MPa and 745 8C (Fig. 5b). The rimcompositions equilibrated at 504 8C and 432 MPa, inagreement with the stability field of retrograde

staurolite and kyanite. The latter is not present in

sample 506 but occurs in samples from a locality

nearby, and is a common retrograde mineral through-

out the garnetcordieritesillimanite zone.

Sample 532A is a migmatite from the garnet

cordierite zone close to the margin of the orthopyr-

oxene zone (Fig. 2), consisting of the assemblage Grt

CrdPlBt1KfsQtz. In sample 532A, biotite1 coex-

ists with garnet and cordierite, suggesting that reaction

(3c) stopped after consumption of sillimanite1. The

garnet zonation (Fig. 7) shows an almandine and

spessartine increase towards the rim whereas pyrope

decreases, leading to a drop in XMg from 0.25 to 0.15.

Using average core compositions of garnet, cordierite,

biotite1, and plagioclase, multiphase equilibria yielded

an intersection of four independent mineral reactions

at 806 8C and 620 MPa (Fig. 5c). A similar pressureof 570590 MPa has been obtained from cordierite

biotitegarnet core compositions (Holdaway and Lee,

1977). We interpret ~800 8C and ~600 MPa asmetamorphic peak conditions along the margin

between the orthopyroxene and the garnetcordier-

itesillimanite zones in the central part of the study

area (Fig. 2). To constrain the PT conditions of the

retrograde equilibration, the TWEEQU 1.02 software

(Berman, 1992) has been used, because it allows the

calculation of mineral reactions involving staurolite.

Average garnet-, cordierite-, biotite-, and plagioclase-

S.H. Buttner et al. / Lrim compositions yielded two intersections at 620 8Cand 456 MPa (three reactions) and 575 8C and 422MPa (four intersections). In good agreement with

these results, conventional biotitegarnet thermometry

(Holdaway et al., 1997; Holdaway, 2000) yielded 590

8C at 400 MPa using average rim compositions ofgarnet and biotite2. Hence, we interpret 420450 MPa

and ~600 8C as conditions of the retrograde equilibra-tion of migmatites in the garnetcordieritesillimanite

zone close to the orthopyroxene zone. The presence of

retrograde sillimanite2 and kyanite (Fig. 3k) indicates

the continuation of the PT path crossing the SilKy

isograd.

Sample 557G is a migmatite of the leucosome

assemblage CrdPlQtzGrt1KfsOpx (Fig. 3e). The

corresponding mesosome consists of biotite1, plagio-

clase, and quartz. Locally, biotite forms a melanosome

along the margin of leucosome and mesosome.

Zoning profiles of cotectic garnet1 are flat, with a

similar composition to garnet cores from sample 532

A (Fig. 7). In contact with biotite of the melanosome,

slightly higher almandine and lower pyrope contents

led to a decrease in XMg from ~0.26 to ~0.21,

suggesting retrograde garnet2 growth and equilibra-

tion. PT values of 812 8C and pressures of 560 MPawere obtained by multiphase equilibria determinations

(Fig. 5d). OpxGrtPl geothermobarometry (Lal,

1993) yielded 825 8C at slightly higher pressure of610 MPa, supported by 810 8C (calculated for 600MPa) obtained from GrtOpx average core composi-

tions (Bhattacharya et al., 1991; Aranovich and

Berman, 1997). Using pairs of single orthopyroxene

and garnet core analyses, temperature determinations

scatter in an interval of 780840 8C (600 MPa). ThePT interval of 560610 MPa and ~800 8C is inagreement with the presence of orthopyroxene in

felsic leucosomes in both fluid-absent and fluid

present conditions (reactions (4a)(4d); see Section

3.4 and Fig. 5a), and is interpreted as the PT

maximum along the eastern margin of the orthopyr-

oxenegarnet zone. Rims of garnet, biotite, plagio-

clase and cordierite are assumed to have re-

equilibrated after the thermal peak, reflecting PT

conditions of the retrograde metamorphism. Average

rim compositions yield an intersection of three

independent mineral reactions at 450 MPa and 705

8C (Fig. 5d). Conventional thermobarometry usinggarnetbiotite thermometry (Holdaway et al., 1997;

83 (2005) 143181 165Holdaway, 2000) and biotitecordierite thermobarom-

etry (Holdaway and Lee, 1977) gives slightly lower

Lithostemperatures but similar pressures of 620 8C and 400430 MPa. We have included the results of both

conventional thermobarometry and multiphase equi-

libria in the interpretation of the retrograde PT path.

Sample MM61 comes from the southwestern part

of the orthopyroxene zone, approximately 6 km west

of the orthopyroxene-in isograd (Fig. 2). This sample

is a coarse-grained, layered diatexite dominated by

QtzKfsPl-rich leucosome, GrtBt2-rich melano-

some, and small relics of a QtzBt1PlIlm-rich

mesosome (Fig. 4g). For geothermobarometry we

used garnet, biotite, and plagioclase core composi-

tions from the neosome. Garnet shows composition-

ally homogeneous cores (Fig. 7). In contact with the

mesosome, biotite and ilmenite inclusions occur

occasionally. Where garnet overgrows the mesosome,

its composition becomes more iron- and calcium-

rich, whereas the pyrope content drops. The spessar-

tine content remains constant, suggesting that the rim

zone reflects retrograde growth (garnet2). Garnet

biotite thermometry using core compositions of

neosome minerals yielded a fairly broad range of

temperatures, depending on the calibration used,

ranging from ~806 8C (Holdaway et al., 1997) to890940 8C (Thompson, 1976; Hodges and Spear,1982) and ~900 8C (TWEEQU 2.02, Berman, 1991;all calculated for 800 MPa). The mineral assemblage

in the neosome includes plagioclase, garnet, and

biotite, but no sillimanite as a magmatic phase.

Sillimanite occurs only as fibrolite along plagioclase

garnet grain boundaries and is most likely a retro-

grade mineral. Hence, the results of GASP barom-

etry, 824 MPa (TWEEQU) and 622 MPa (Holdaway,

2000), are both controversial and equivocal. How-

ever, the intersection of the multiphase equilibria

curves in particular (~900 8C and 824 MPa; Fig. 5e)is in good agreement with the metamorphic field

gradient defined by peak conditions of samples 513,

506, 532A, and 557G (Fig. 5a). Garnetbiotite

temperatures of ~900 8C are lower than the biotitebreakdown at low aH2O (~950 8C/800 MPa;Vielzeuf and Montel, 1994), and therefore not

unrealistic for sample MM61, which comes from

the highest-grade metamorphic part of the study area.

Using average biotite, garnet and plagioclase rim

compositions, garnetbiotite thermometry and GASP

S.H. Buttner et al. /166yield 595 MPa and 694 8C, using TWEEQU 2.02.(Berman, 1991). Again, garnetbiotite temperaturesand GASP pressures using the calibrations of Hold-

away et al. (1997) and Holdaway (2000, 2001) are

lower in P and T (386 MPa and 652 8C for averagerim compositions). The results of geothermobarom-

etry from sample MM61 confirm the high peak

temperatures, exceeding 800 8C at middle to lowercrustal levels in the orthopyroxene zone, which have

been obtained from sample 557G and from petroge-

netic grids. The pressures calculated with conven-

tional methods and multiphase equilibria differ by

~200 MPa for both, peak and retrograde equilibra-

tions. The higher PT values from multiphase

equilibria appear more plausible because the migma-

tites at the sample location MM61 should have

experienced higher pressure and temperature than the

samples from close to the orthopyroxene-in isograd

(532A and 557G). For methodological reasons

(uncertain presence of Sil at the thermal peak) we

have not used the results from MM61 to support our

geodynamic model (see below).

Sample 542 is a medium-grained mylonite from a

retrograde shear zone within the orthopyroxene zone,

consisting of the assemblage QtzBtGrtSilPl

pinite. Elongated quartz lenses show chessboard

patterns indicating deformation and annealing within

the stability field of high-quartz (Kruhl, 1996). Both

multiphase equilibria (Fig. 5f) and conventional

garnetbiotite/GASP thermobarometry (TWEEQU

2.02; Holdaway et al., 1997; Holdaway, 2000,

2001) yield nearly identical results for both, core

and rim compositions. Average core compositions

yield 430 MPa and 665 8C (conventional methods)and 449 MPa and 675 8C with a slightly higher GrtBt temperature estimation (TWEEQU 2.02; Fig. 5f).

These conditions were interpreted as the PT

minimum for the onset of retrograde plastic solid

state shearing in the orthopyroxene zone. Using the

same methods as for core compositions, the rim

compositions equilibrated at 440 MPa/610 8C and436 MPa/612 8C, respectively. These PT dataprobably reflect post-kinematic equilibration, because

chessboard subgrain patterns in quartz indicate the

formation of the quartz lenses at higher temperatures,

in the stability field of high-quartz (Kruhl, 1996). As

sample 542 comes from the orthopyroxene zone, it is

evident that the peak conditions of the pre-mylonitic

83 (2005) 143181protolith were similar to those of samples 557G or

MM61.

5. Geochronology

In order to gather age data on regional meta-

morphism, magmatic activity, retrograde shearing,

and subsequent cooling, we selected six felsic

pegmatites, one garnetaplite, one tonalite, two

migmatites, two low-grade metapelites, and two

calcsilicate rocks for isotopic dating. The pegmatites

have been described in Section 3.6, while petro-

graphic descriptions of other dated rock types can be

on Re-single filaments. U and Pb isotopic analyses

were conducted using a Finnigan MAT 262 at the

GeoForschungsZentrum (GFZ) Postdam. Instrumental

mass-fractionation was corrected by 0.1% per a.m.u.

The 2r reproducibility of the NBS SRM 981 Pbstandard is better than 0.1% for the 206Pb/204Pb and207Pb/204Pb ratios. Monazite U/Pb data were corrected

for excess 206Pb following Scharer (1984), assuming a

whole rock Th/U ratio of 0.5. The raw data were

processed and ages calculated using the program

] STP

S.H. Buttner et al. / Lithos 83 (2005) 143181 167found in the Electronic Data Supplement. Sample

locations are shown in Fig. 2. KAr and 40Ar39Ar

data on muscovite are presented in Table 2 and Fig. 8.

RbSr, SmNd, and UPb internal mineral system-

atics are given in Tables 35, and in the Electronic

Data Supplement. The geochronological data set is

summarised in Table 6. Petrographic descriptions and

additional information on the geochonological data set

can be found in the Electronic Data Supplement. The

critical mineral assemblage of each dated rock is

given in Table 6.

5.1. Analytical methods

5.1.1. UPb

Monazite and titanite were separated using a

magnetic separator and heavy liquids. Grains free of

inclusions and alterations were hand-picked under the

binocular microscope. Two different size fractions of

monazite (~200 Am and ~100 Am) could be distin-guished in the sample 465. Before the addition of a205Pb235U tracer and dissolution of the sample in

conc. H2SO4 (monazite) or HF (titanite) on a hot

plate, the grains were cleaned in hot, dilute HNO3,

ultraclean H2O, and acetone. Standard methods were

used for chemical separation of Pb and U and loading

Table 2

KAr age determination

Sample Spike [no.] K2O [wt.%]40Ar* [nl/g

472 B Ms 2247 10.68 162.91

515 B Ms 2246 10.23 171.35

552 Ms 2245 10.21 161.87

526 Ms 2244 10.41 171.29

620 A Msb2Am 2699 5.69 83.08620 A Ms b0.2Am 2701 5.35 75.57

620 C Ms b2Am 2700 6.29 90.72620 C Ms b0.2Am 2693 6.08 82.69packages PBDAT (rev. 1.24; Ludwig, 1993) and

Isoplot/Ex (rev. 2.49; Ludwig, 2001).

5.1.2. SmNd

Sm and Nd concentrations were determined by

isotope dilution using a mixed 149Sm150Nd spike.

Sm and Nd isotope analyses were carried out on a

Finnigan MAT 262 at the GFZ Postdam. Nd was

analysed in dynamic, Sm in static mode. For age

calculations, standard errors of F0.004% for143Nd/144Nd ratios and F1% for 147Sm/144Nd ratioswere assigned to the results. The value obtained for143Nd/144Nd of the La Jolla standard during the period

of analytical work was 0.511852F0.000006 (n=6).Further analytical details are given in Kuhn et al.

(2001).

5.1.3. RbSr

Both Rb and Sr concentrations were determined by

isotope dilution using mixed 87Rb84Sr spikes.

Determinations of Sr isotope ratios were carried out

on a VG Sector 54 multicollector TIMS instrument

(GFZ Potsdam) in dynamic mode. The value obtained

for 87Sr/86Sr of the NBS standard SRM 987 during the

period of analytical work was 0.710266F0.000012(n=25). Rb analyses were done on a VG Isomass 54

40Ar* [%] Age [Ma] 2r-Error [Ma] 2r-Error [%]

96.33 420 12 2.8

97.75 457 14 2.9

98.31 435 10 2.3

97.00 450 11 2.4

98.49 404 8 2.0

97.51 392 8 2.197.99 400 8 2.0

96.27 379 8 2.1

00

00

Lithos300

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400

450

500

300

350

400

450

500

0 20 40 60 80 1

app

aren

t age

[Ma]

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t age

[Ma]

472 B

515 B

total gas age: 408 7 Ma

total gas age: 454 8 Ma

cumulative percentage 39Ar released

0 20 40 60 80 139

S.H. Buttner et al. /168single collector mass spectrometer (GFZ Postdam).

The observed ratios were corrected for 0.25% per

a.m.u. mass fractionation. Total procedural blanks

were consistently below 0.15 ng for both Rb and Sr.

For calculations of isochron parameters, standard

errors of F0.005% for 87Sr/86Sr ratios and of F1%for Rb/Sr ratios were applied if individual analytical

errors were smaller than these values. Further details

are given in Hetzel and Glodny (2002).

5.1.4. KAr

Prior to analysis, purified micas were ground in

pure ethanol to remove altered rims that might have

suffered a loss of Ar or K. Argon isotopic composi-

tions were measured in a Pyrex glass extraction and

purification line coupled to a VG 1200 C noble gas

mass spectrometer operating in static mode. The

amount of radiogenic 40Ar was determined by isotope

dilution using a highly enriched 38Ar spike from

Schumacher, Bern (Schumacher, 1975). The spike

was calibrated against the biotite standard HD-B1

(Fuhrmann et al., 1987). The age calculations are

based on the radioactive decay constants recommen-

cumulative percentage Ar released

Fig. 8. 40Ar39Ar laser step heating age-spectra of pegmatitic muscovite.

symbols represents F1r.MM 110

MM 72 B

integrated age: 408 7 Ma

integrated age: 442 8 Ma

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