The Proto-Andean Margin of Gondwana (Geological Society Special Publication No. 142)

The Proto-Andean Margin of Gondwana Geological Society Special Publications Series Editors A. J. FLEET A. C. MORTON A...

Author: R. J. Pankhurst | C. W. Rapela

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The Proto-Andean Margin of Gondwana

Geological Society Special Publications Series Editors A. J. FLEET A. C. MORTON A. M. ROBERTS

It is recommended that reference to all or part of this book should be made in one of the following ways. KELLER, M., BUGGISCH, W. & LEHNERT, O. 1998. The stratigraphical record of the Argentine Precordillera and its plate-tectonic background. In: PANKHURST, R. J. & RAPELA, C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 35-56. PANKHURST, R. J. & RAPELA, C. W. (eds) 1998. The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 142

The Proto-Andean Margin of Gondwana

EDITED BY

R O B E R T J. P A N K H U R S T British Antarctic Survey and NERC Isotope Geosciences Laboratory, UK AND

CARLOS W. R A P E L A Universidad Nacional de La Plata, Argentina

1998 Published by The Geological Society London

THE GEOLOGICAL SOCIETY The Society was founded in 1807 as The Geological Society of London and is the oldest geological society in the world. It received its Royal Charter in 1825 for the purpose of 'investigating the mineral structure of the Earth'. The Society is Britain's national society for geology with a membership of around 8500. It has countrywide coverage and approximately 1500 members reside overseas. The Society is responsible for all aspects of the geological sciences including professional matters. The Society has its own publishing house, which produces the Society's international journals, books and maps, and which acts as the European distributor for publications of the American Association of Petroleum Geologists, SEPM and the Geological Society of America. Fellowship is open to those holding a recognized honours degree in geology or cognate subject and who have at least two years' relevant postgraduate experience, or who have not less than six years' relevant experience in geology or a cognate subject. A Fellow who has not less than five years' relevant postgraduate experience in the practice of geology may apply for validation and, subject to approval, may be able to use the designatory letters C Geol (Chartered Geologist). Further information about the Society is available from the Membership Manager, The Geological Society, Burlington House, Piccadilly, London WlV 0JU, UK. The Society is a Registered Charity, No. 210161. Published by The Geological Society from: The Geological Society Publishing House Unit 7, Brassmill Enterprise Centre Brassmill Lane Bath BA1 3JN UK (Orders: Tel. 01225 445046 Fax 01225 442836)

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Contents

PANKHURST, R. J. & RAPELA, C. W. The proto-Andean margin of Gondwana: an introduction

1

ASTINI, R. A. Stratigraphical evidence supporting the rifting, drifting and collision of the Laurentian Precordillera terrane of western Argentina

11

KELLER, M., BUGGISCH, W. & LEHNERT, O. The stratigraphical record of the Argentine Precordillera and its plate-tectonic background

35

BENEDETTO, J. L. Early Palaeozoic brachiopods and associated shelly faunas from western Gondwana: their bearing on the geodynamic history of the pre-Andean margin

57

DICKERSON, P. W. & KELLER, M. The Argentine Precordillera: its odyssey from the Laurentian Ouachita margin towards the Sierras Pampeanas of Gondwana

85

HUFF, W. D., BERGSTRC)M, S. M., KOLATA, D. R., CINGOLANI, C. A. ~; ASTINI, R . A . Ordovician K-bentonites in the Argentine Precordillera: relations to Gondwana margin evolution

107

BAHLBURG, H. The geochemistry and provenance of Ordovician turbidites in the Argentine Puna

127

RAMOS, V. A., DALLMEYER, R. D. & VUJOVICH, G. I. Time constraints on the Early Palaeozoic docking of the Precordillera, central Argentina

143

VUJOVICH, G. I. & KAY, S. M. A Laurentian? Grenville-age oceanic arc/back-arc terrane in the Sierra de Pie de Palo, Western Sierras Pampeanas, Argentina

159

RAPELA, C. W., PANKHURST, R. J., CASQUET,C., BALDO, E., SAAVEDRA,J., GALINDO, C. & FANNING, C. M. The Pampean Orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de C6rdoba

181

DALLA SALDA,L. H., LOPEZDE LUCHI, M. G., CINGOLANI, C. A. & VARELA, R. LaurentiaGondwana collision: the origin of the Famatinian-Appalachian orogenic belt (a review)

219

VON GOSEN, W. & PROZZI, C. Structural evolution of the Sierra de San Luis (Eastern Sierras Pampeanas, Argentina): implications for the proto-Andean margin of Gondwana

235

SIMS, J. P., IRELAND, T. R., CAMACHO, A., LYONS, P., PIETERS, P. E., SKIRROW, R.G., STUART-SMITH, P. G. • MIRO R. U-Pb, Th-Pb and Ar-Ar geochronology from the southern Sierras Pampeanas, Argentina: implications for the Palaeozoic tectonic evolution of the western Gondwana margin

259

SAAVEDRA, J., TOSELLI, A., ROSSI, J., PELLITERO, E. & DURAND, F. The Early Palaeozoic magmatic record of the Famatina System: a review

283

GRISSOM, G. C., DEBARI, S. M. & SNEE, L. W. Geology of the Sierra de Fiambalfi, northwest Argentina: implications for Early Palaeozoic Andean tectonics LLAMBiAS, E. J., SATO, A. M., ORTIZ SUAREZ, A. & PROZZI, C. The granitoids of the Sierra de San Luis

297

PANKHURST, R. J., RAPELA, C. W., SAAVEDRA,J., BALDO, E., DAHLQUIST,J., PASCUA, I. & FANNING, C. M. The Famatinian magmatic arc in the central Sierras Pampeanas: an Early to Mid-Ordovician continental arc on the Gondwana margin

343

Index

369

325

The proto-Andean margin of Gondwana: an introduction R. J. P A N K H U R S T 1 & C. W. R A P E L A 2

1 British Antarctic Survey, c/o N E R C Isotope Geosciences Laboratory, Kingsley Dunham Centre, Keyworth, Nottingham NG12 5GG, UK 2 Centro de Investigaciones Geol6gicas, Universidad Naeional de la Plata, 644 Calle No. 1, 1900 La Plata, Argentina Abstract: A basic background is presented for the discussion of the Early Palaeozoic

geology of western Argentina covered by this book. This includes the definition and terminology of orogenic cycles on this part of the Gondwana margin, represented by the Eastern Sierras Pampeanas. The Pampean orogeny (Early Cambrian) relates to an intense but short-lived period of terrane collision predating the rifting of the Precordillera terrane from Laurentia. The Famatinian cycle is predominantly represented by intense subductionrelated magmatism of Early-Middle Ordovician age, developed on the continental margin of Gondwana during the rifting and drifting of the Precordillera terrane. The Grenvillian basement of the latter is further exemplified by a new Rb-Sr whole-rock isochron age of 1021 + 12Ma for orthogneisses from the Western Sierras Pampeanas. A mid-Ordovician granite in this area (dated at 481 4-6 Ma by U-Pb ion microprobe data) may be related to rifting while the Precordillera terrane was still attached to Laurentia. A divergence of opinion is pointed out between some authors in this book who favour mid-Ordovician collision of the Precordillera with Gondwana, and others who place it much latter, in Silurian or Devonian times. The Palaeozoic geological history of western South America has been the focus of intense international interest during the last five years. The development of this subject, and the growing state of our knowledge, can be traced from the idea of Palaeozoic continent-continent collision between North and South America along this margin (Dalla Salda et al. 1992a, b), through the proceedings of international conferences in M6rida, Spain ('El Paleozoico Inferior de IberoAmerica', Gutierrez-Marco & Saavedra 1992), in San Juan, Argentina (Penrose Conference 'The Argentine Precordillera: a Laurentian Terrane', Dalziel et al. 1996), and the international symposium 'The Proto-Andean Margin of Gondwana', co-organized by IGCP projects 345 and 376 (XIII Congreso Geol6gico Argentino, Actas V, 257-561, Buenos Aires, 1996). A major landmark was the overall consensus reached by participants in the 1995 Penrose Conference that the Precordillera of western Argentina is an allochthonous terrane, originally detached from Laurentia (Dalziel 1997 and references therein; see also Dickerson & Keller this volume) and accreted to the southwestern margin of Gondwana. The timing of this transfer, and the events and processes involved in the arrival and collision of the Precordillera allochthon, have far-reaching consequences for geological, biological and palaeogeographical interpretations and the distribution of mineral resources. Many of the implications are only just beginning to be understood.

The papers presented in this book come mainly from participants in the Buenos Aires meeting in October 1996, where the latest results from various Palaeozoic provinces were discussed, including recent information on the history of the Gondwana foreland before and during the accretion of the Precordillera terrane. For a general audience, appreciation of the arguments presented requires a basic knowledge of the common geographical, geological and tectonic terminology used by South American geologists. It is the purpose of this introduction to provide an adequate background for the non-specialist. The region concerned and its relevant geographical features are illustrated in Figs 1 and 2. The Sierras Pampeanas of western Argentina are elevated ranges of pre-Cenozoic basement rock, thought to result from uplift on reverse faults during upper Tertiary Andean deformation. This structural domain roughly coincides with a near-horizontal segment of the presently subducting Nazca p l a t e - the 'flat-slab zone' between 27~ t and 33~ (Isacks 1988; Jordan & Allmendinger 1986) (Fig. 1), also characterized by Miocene arc magmatism more than 700km inboard of the Chile trench (Kay & Gordillo 1994). The pre-Mesozoic geological record of the Sierras Pampeanas and the immediately adjacent regions is characterized by evidence of several periods of convergence along the proto-Pacific margin of Gondwana. This is the basis of the recognition of the major orogenic cycles, such as

PANKHURST, R. J. & RAPELA, C. W. 1998. The proto-Andean margin of Gondwana: an introduction. In: PANKHURST,R. J. & RAPELA, C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, St3eciai Publications, 142, 1-9.

2

R. J. PANKHURST & C. W. RAPELA

Fig. 1. Sketch map of southern South America, showing the location of the main areas of Lower Palaeozoic outcrop in western Argentina and their general nomenclature.

the Pampean and Famatinian (Acefiolaza & Toselli 1973), Gondwanian (Llambias et al. 1984) and Andean cycles; a brief description of the meaning that we attach to each of these terms is given in Table 1. Based on local and distinct geological characteristics, they have effectively displaced imported European and North American terms such as 'Caledonian', 'Hercynian', 'Acadian' and 'Taconic', although ironically, with a new and better understanding of continent and supercontinent configurations, the real equivalence of some of these events is now ripe for re-evaluation. Recognition, definition and ordering of the discrete phases of deformation, metamorphism and magmatism that make up these cycles requires intensive research and interpretation, and is thus sometimes a contentious subject. The chronostratigraphical limits of these large-scale geological cycles in southern South America can be ascribed to major plate re-arrangements in the evolutionary stages of the Gondwana supercontinent. The changes between these stages, and therefore in the geodynamic environment along the margin, are mostly marked by initiation, or sudden acceleration, of subduction (see Table 1). It must be stressed that the word cycle

Fig. 2. Sketch map of the region of the Sierras Pampeanas and the Argentine Precordillera, the area encompassing the subject of most papers in this volume.

INTRODUCTION

3

Table 1. Phanerozoic geological cycles along the Pacific margin of South America

Pampean (Neoproterozoic-Late Cambrian) First recognized in the basement of the Eastern Cordillera, the Puna and the northwestern Sierras Pampeanas (Acefiolaza et al. 1990 and references therein), this Neoproterozoic-Cambrian cycle is now regarded as a major orogenic event in southwestern Gondwana. Along the southern proto-Andean margin, the Pampean cycle includes formation of a passive margin sedimentary basin, emplacement of metaluminous granitoid sequences, penetrative folding and low-high-grade metamorphism during mid-Cambrian time (see Rapela et al. this volume), and extensional collapse in the late Cambrian. The Pampean orogeny is broadly coeval with the collisional Rio Doce Orogeny, the last event of the Brasiliano-Pan African cycle along the southern coastal region of Brazil (Campos Neto & Figueiredo 1995). Famatinian (Early Ordovician-Early Carboniferous) A major accretional and orogenic episode that started with subduction along the Cambrian proto-Pacific margin (at c. 490 Ma). The Famatinian orogeny in the Gondwana foreland and the collision of the Precordilleran Terrane along the continental margin are usually both ascribed to this cycle, although they may not be causally related (see Pankhurst et al. this volume). Most of the contributions in this book are devoted to different aspects of this complex process, a summary of which is presented in Fig. 3. It should be noted, however, that mid-Palaeozoic tectonic phases registered in the clastic sequences of the Precordillera ('Precordilleran' and 'Chanic' tectonic phases, Fig. 3) have been related to the subsequent collision of another terrane ('Chilenia') to the west of the Precordillera (Ramos et al. 1986; Astini 1996). These tectonic events, as well as the broadly coeval intrusion of within-plate granites in the Gondwana foreland, are here tentatively included as the latest events of the Famatinian Cycle (see Fig. 3). Gondwanian (Early Carboniferous-Early Cretaceous) This was a period of the maximum extent and relative stability of the super-continent, from Lower Carboniferous (c. 330-340 Ma, see insert, Fig. 3) to final break-up. Following the Early and mid-Palaeozoic orogenies and accretionary events, intrusion of Carboniferous Cordilleran-type batholiths associated with a new subduction regime along the palaeo-Pacific margin heralded a major plate re-assembly. Development of extensive Permian to Jurassic rhyolitic provinces and inner cordilleran plutonic belts are characteristic of the Gondwanian cycle (Rapela et al. 1996). Andean (Early Cretaceous-Present) From the final break-up of Gondwana to the present. The Andean Cycle started with the opening of the South Atlantic and separation of South America from the African-Indian plate. This event was closely associated with extrusion of the Paranh-Etendeka flood basalt province (c. 130 Ma Renne et al. 1992; Turner et al. 1994). The generation of plutonic and volcanic rocks typical of the modern Andes became predominant as the convergence velocity at the palaeo-Pacific margin of the newly formed South American plate increased. The term 'Andean Orogeny' is linked to several compressive phases during this period, but is most commonly used for those of Tertiary and modern age. Formation of intra-continental extensional basins and a passive Atlantic margin are also typical of the Andean period.

is used here to describe stages during the accretion and break-up of a supercontinent, and not necessarily a regular or repeated succession of events. Not even the magmatic products of these subduction regimes are identical; the Famatinian, Gondwanian and Andean calcalkaline rocks each have their own distinctive characteristics. A summary of the major geological features and inferred tectonic events for the Precambrian and Early to mid-Palaeozoic provinces from 22 to 33~ is shown in Fig. 3, based on established literature but in some cases modified by the new results presented in this volume. The basic distinction of a Precordillera terrane and a Gondwana foreland seems useful to describe and contrast the main features of the adjacent continental masses that are inferred to have been amalgamated during the Lower Palaeozoic. The term Precordillera terrane is used to encompass

the Lower Palaeozoic sedimentary sequences of the Precordillera, and its Grenvillian-age basement represented by the adjacent parts of the Western Sierras Pampeanas, as in Astini et al. (1995); see also Ramos et al. (this volume) who use the name 'Cuyania' for this terrane. Early Palaeozoic magmatism and metamorphism is clearly recorded in central and N W Argentina (Sierras Pampeanas, Sierra de Famatina and the Puna), extending into southern Bolivia, northeastern Chile, and the coastal areas of southwestern Peru. The main components are the Ordovician plutonic and volcanic rocks of the Puna, the Famatina System, and the Central Batholithic Belt, which together constitute the Famatinian magmatic arc, and the Cambrian metamorphic and plutonic igneous sequences of the Eastern Sierras Pampeanas and the Eastern Cordillera (Fig. 1). Sedimentation in the Precordillera and granitoid magmatism in

4

R.J.

PANKHURST

& C. W .

RAPELA

O R O G E N I C E V E N T S IN T H E P R O T O - P A C I F I C M A R G I N OF G O N D W A N A System and

Gondwana

dillera T e r r a n e

foreland

Series I Pampean foreland :

[PC (~)] and

The Famatinian arc :

ts Pampeanas [WSP (~)1

Puna [P (~)], Famatina System (FS (~)] & Central batholithic belt [CBB (~)]

Eastern Cordillera [EC (~)1 & Eastern Sierras Pampeanas [ESP (~)]

of cordilleran plutons: CARBONIFEROU;

anicextensive tectonic phase mic basic and ultrabasic rocks layered with Silurian?)nian sediments

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- ?Continued magmatism, including emplacement of post-compressional monzogranite plutons - Syntectonic metaluminous batholiths. Calc-alkaline volcanics interlayered with shallow marine volcaniclastics. Opening of backarc basin.

- Transpressive structures associated with low-to-high grade metamorphism. Inner cordilleran plutons. High-A1 and Na trondhjemitic bodies. A

Famatinian Cycle i i i

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] PRECAMBRIAN

between Lower Ordovician & Ashgillian deposits in EC and P. Regional mylonites in the ESP.

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i i

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Intrusion of large post-orogenic batholiths dominated by high-K porphyritic monzogranitic facies in the ESP (e.g. Achaia batholith) and the CBB (e.g. Capillitas and Velasco batholiths). Cordierite -bearing facies are common in the CBB

the PC Caradoc

Cycle

Intrusion of within- plate REE-enriched plutons

oic "Grenville" metamorphic the WSP and basement of the lites, marbles, schists and ing immature arc or oceanic ~s.

- Extension associated with collapse of the Pampean orogen. Late Cambrian platform sediments unconformably overlie the folded Puncoviscana Fm (EC, P). Alkalic volcanism in the EC & P. - Obduction of MORB-type ophiolites in the ESP. - Early-to-Mid-Cambrian compressive events. Folding and low-grade (EC, FS) to high- grade (ESP) metamorphism of the Puncoviscana Fm. and equivalents. Widespread anatexis and S-type plutonism in the ESE - Intrusion of metaluminous orthogneisses. - Passive margin sequence composed of widespread thick turbiditic deposits carrying Neoproterozoic to Lower Cambrian trace fossils in the EC and eastern Puna; southern equivalents in the FS, CBB and ESP, the Puncoviscana Formation.

Pampean Cycle Arequipa-Antofalla craton IAA (~)] - Early Proterozoic igneous and sedimentary rocks with an overprinted high-grade Grenville age event in the late Middle Proterozoic. Ages ranging from 1900 to 980 My.

Fig. 3. S c h e m a t i c r e p r e s e n t a t i o n o f the m a i n g e o l o g i c a l events r e c o g n i z e d in the P a l a e o z o i c h i s t o r y o f the region, with a c o m p a r i s o n o f s u c h event in the a l l o c h t h o n o u s P r e c o r d i l l e r a t e r r a n e a n d the G o n d w a n a f o r e l a n d ( E a s t e r n a n d W e s t e r n Sierras P a m p e a n a s ) .

INTRODUCTION the Eastern Sierras Pampeanas continued into Devonian-Carboniferous times, after the final assembly of this accreted margin. Carboniferous calc-alkaline magmatism to the west of the Precordillera, in Argentina and Chile, marks a yet later stage in the development of the Andean active margin that is not considered here. The uplifted Palaeozoic rocks of the foreland reach altitudes of over 6000 m in the Famatina System. The level of crustal exposure seems to be higher in the north, with shallow marine sediments of Lower Ordovician age and coeval volcanic and volcaniclastic sequences in the Famatina System and the Puna (Saavedra et al. this volume), and Neoproterozoic-Lower Cambrian turbiditic sediments with trace fossils in the Eastern Cordillera (see review by Bahlburg & Herv~ 1997; Bahlburg this volume). In contrast, such supracrustal rocks are relatively rare in the southern areas, where both low- and medium-tohigh grade metamorphic rocks predominate as the host rocks to Ordovician magmatism. South of 33~ the typical uplifted foreland block of the 'fiat-slab' segment abruptly disappears. However, rocks carrying fossils typical of the carbonate platform of the Precordillera have been described at 35~ 68~ in the Ponon Trehue area (Bordonaro et al. 1996), and poorly preserved carbonate rocks of La Pampa at 37~ 67~ are also considered to be a southeasterly extension of the Precordilleran sequences (Criado Roqu6 1972; Astini et al. 1995). Scattered outcrops of igneous and metamorphic rocks in this region have given late Proterozoic K-Ar ages (Linares et al. 1980), suggesting that the crystalline basement of the Western Sierras Pampeanas might also extend to these southerly latitudes (see Fig. 1 and Ramos et al. this volume). Equivalents of the CambrianLower Ordovician carbonate platform of the Precordillera, the key sequences that demonstrate a Laurentian biogeographical affinity (Benedetto this volume and references therein), do not crop out either in Patagonia or to the north of the Precordillera; whether this is a real geological discontinuity or a coincidence of tectonic or erosion processes is still not well understood but is clearly important to the terrane accretion concept (see below).

The Precordillera Terrane Although the basement of the Precordillera is not exposed, indirect evidence of its 'Grenvillian' characteristics and age come from U-Pb studies of xenoliths from Miocene lava domes (Kay et al. 1996). This has been taken as further support for

5

the hypothesis that the Precordillera was derived from part of Laurentia corresponding to the southern Appalachians (e.g., Dalla Salda et al. 1992a; Astini et al. 1995). Other 'Grenvillian ages' (c. 1000 i 100 Ma) have been obtained in the adjacent Sierra de Pie de Palo to the east, including U-Pb ages on orthogneiss, diorite, rhyolite and amphibolite (938 to 1091Ma, McDonough et al. 1993, analytical data and errors not reported). This has lead to a growing consensus that the whole of the Western Sierras Pampeanas might represent basement remnants of an allochthonous Grenvillian terrane of which the Precordillera is only a part. Rb-Sr analysis of equivalent orthogneiss suites in the Sierra de Umango, 250km to the north of Pie de Palo (Fig. 2), has produced a similar age of 1030 + 30Ma (Varela et al. 1996). Finally, we have ourselves obtained a new Rb-Sr whole-rock isochron of 1021 + 12 Ma (2a; MSWD = 2.9) for seven samples of granitic orthogneisses from the highest part of the Sierra de Pie de Palo (Fig. 4). Together with the initial 87Sr/86Sr ratio of 0.7045 +0.0003, this shows that the parent magmas of these gneisses were generated in Grenvillian times from an immature source. Significantly, in view of the intense Palaeozoic metamorphic and magmatic history prevalent to the east, their Rb-Sr systems show no evidence of subsequent events. The evidence for possible Laurentian derivation of the Precambrian metamorphic sequences to the north, west, and south of the Sierras Pampeanas, is more difficult to evaluate. Small outcrops of pre-Carboniferous metamorphic rocks west of the Precordillera in both Argentina and Chile have been taken as evidence for a later exotic terrane ('Chilenia', Ramos et al. 1986), and some of these have now yielded Grenvilleage zircon (Ramos & Basei 1997). Their reality as a separate terrane, or as part of the basement of the Precordillera terrane, is not clear, and the extent to which more distant metamorphic terfanes along the proto-Andean margin can be assigned 'Laurentian' derivation is largely model dependent. For example, the Laurentian 'Occidentalia Terrane' of Dalla Salda et al. (1992a, b, and this volume), includes all known Precambrian outcrops along the Andean margin: the Arequipa-Antofalla craton, the Western Sierras Pampeanas and the western edge of the Patagonian massifs. A continent-continent collision between eastern Laurentia and southwestern Gondwana along the whole length of this margin is postulated in this model. However, Dalla Salda et aL (this volume) base many correlations on the original, very broad, definitions of the Pampean and Famatinian cycles (Acefiolaza &

6

R. J. PANKHURST & C. W. RAPELA

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,

12

Rb-Sr analytical data for orthogneisses and migmatites, Sierra de Pie de Palo: Sample Rb ppm Sr ppm 87Rb/86Sr % s . d . 87Sr/86Sr SPP407 Gneiss 53.2 74.7 2.0639 0.5 0.734410 SPP411 Gneiss 61.8 116.5 1.5384 0.5 0.727311 SPP412 Gneiss 44.8 131.3 0.9891 0.6 0.719007 SPP413 Granite 96.5 90.8 3.0853 0.5 0.749937 SPP414A Gneiss 41.1 121.9 0.9755 0.7 0.718620 SPP414B Granite 88.3 104.5 2.4525 0.5 0.740044 SPP416 Gneiss 27.9 100.4 0.8036 1.0 0.716315 Sample localities: SPP 407, Quebrada La Petaquita (31~ 68~ remainder around antenna (31~ 67~ 8ZSr/SSSrprecise to :1:0.01%, relative to NBS987= 0.710235

,

,

I

,

14

television

Fig. 4. New geochronological data relating to the Sierra de Pie de Palo, Western Sierras Pampeanas. (a) Rb-Sr isochron for orthogneisses from the central (highest) part of the sierra, indicating a Grenvillian age of formation, (b) U-Pb SHRIMP data for a granite emplaced into metamorphic basement in the SE part of the sierra, corresponding to an Early to Mid-Ordovician age (data available from R.J.P.).

Toselli 1973), which are at variance with the more restricted definitions shown in Fig. 3 and adopted by other authors in this volume (Pankburst et aL; Rapela et aL; Sims et aL). Moreover, Pb isotope studies in the Arequipa-Antofalla craton have been taken to suggest that it represents a crustal reservoir distinct from that of the Western Sierras Pampeanas and the Precordillera basement (Tosdal 1996). Another approach is that of Astini et al. (1995), Thomas & Astini (1996), and Astini (this volume), who consider the platform of the Precordillera and its Grenvillian basement as a more restricted microcontinent (the Precordillera Terrane or Cuyania), detached from Laurentia during the Lower Cambrian and accreted to Gondwana in mid-Ordovician times after travelling independently across the Iapetus Ocean. The size of this terrane has been estimated as about 1000km long and 400kin wide when Andean shortening is restored (Keller & Dickerson 1996), or as a block c. 800km square (Thomas & Astini 1996). Dalziel (1997) proposed an alternative model in which the micro-continent that collided with Gondwana

was a stretched passive margin plateau that was still connected to Laurentia during a MidOrdovician collision. Doubt is cast on this model by current considerations of stratigraphical and palaeontological evidence (Asfini, this volume; Benedetto this volume). New ages and petrology relating the Grenville basement of the Western Sierras Pampeanas and its deformation are presented by Ramos et al. (this volume). The geochemistry and tectonic interpretation of mafic sequences in these western areas as representing a variety of Precambrian oceanic environments are considered by Vujovich & Kay (this volume) who identifiy some of the most primitive Precambrian magma types known. Also shown in Fig. 4 are new U - P b zircon data obtained by SHRIMP (ANU, Canberra) for a granite cutting the metamorphic sequence, Zircon cores show a range of inherited ages up to at least 1400Ma, whereas data from the rims indicate an emplacement age of 481 i 6 Ma. If this were considered part of the Famatinian cycle, it would imply that the Grenvillian basement of the Sierra de Pie de Palo had already collided with the Gondwana foreland by Early

INTRODUCTION Ordovician times. This would be consistent with the models of Astini and Ramos et al. (this volume), who argue for mid-Ordovician collision. However, models in which the collision of the Precordillera terrane occurred much later, in Silurian-Devonian times, are supported by some of the new evidence and interpretations presented in this volume (see Keller et al.; Benedetto; Pankhurst e t ai.). The palaeontological content of the PrecordiUera is revisited by Benedetto, stressing the palaeo-biogeographical links with eastern Laurentia. He suggests that the Precordillera and Gondwana faunas show differences throughout the Late Ordovician and that the first unequivocal evidence for geographical continuity between them seems to be in the Wenlockian. Keller et al. (this volume) identify major unconformities in the Precordillera succession as relating to the rifting and drifting stages of separation from Laurentia, and refute the possibility of an Ordovician collision with Gondwana. Von Gosen & Prozzi (this volume) consider that post-Famatinian compression in the Sierra de San Luis was probably synchronous with collision. The La Escalerilla granite, affected by this deformation, is of Devonian age according to Stuart-Smith et al. (unpublished, cited by Sims et al. this volume). Finally, the subduction episode that closed the Southern Iapetus Ocean, resulting in the collision of the Precordillera terrane, did not start until c.490Ma (Pankhurst e t aL this volume; Sims et al. this volume) and it is therefore likely that the collision itself was much later. Our suggestion is that the mid-Ordovician granite in the Sierra de Pie de Palo is related to magmatism during the initial rifting of the terrane from Laurentia rather than to Gondwana events. Mid-Ordovician K-bentonites intercalated in the Precordillera carbonate sequences might be explained in the same way, although some authors in the present volume (Huff et al.; Dalla Salda et al.) prefer their previous interpretation of derivation from the Famatinian arc itself during a mid-Ordovician collisional episode.

The Gondwana foreland Most of the geological constraints for the collision models outlined above have come from recent sedimentological and biostratigraphical studies on the Precordilleran sequences and petrological-geochronological studies of its Grenvillian basement. It is nonetheless obvious from Fig. 3 that some of these models will be severely constrained, or even discarded, when

7

the timing, geochemical signatures and P - T - t paths of the foreland sequences and the Famatinian arc can be compared in detail with those of the inferred exotic terranes. One purpose of this book is to initiate progress towards that goal, as well as to encourage further research on the many large igneous-metamorphic areas that still lack detailed structural, petrological and isotopic studies. Rapela e t al. and Sims et al. discuss the Cambrian to Lower Ordovician geological characteristics of the Gondwana continental margin before the collision of the Precordillera terrane, represented by the Eastern Sierras Pampeanas and the Eastern Cordillera. Based on a new structural, metamorphic, geochemical, and geochronological study, Rapela et al. conclude that the Pampean orogeny itself resulted from a major earlier collisional event in Early to MidCambrian times.

The Famatinian arc Finally, the development of the Famatinian arc and subsequent transformation of the margin into the Gondwana foreland is treated in number of papers in this volume; in the Famatina System (Saavedra et al.), the Sierra de Fi~tmbala (Grissom et al.), the central southern Sierras Pampeanas (Pankhurst et al.), and the Sierras de San Luis (Llambias et al.; yon Gosen et aL). Further evidence for the geochronological constraints on the timing of both Pampean and Famatinian orogenic events is provided by Sims et al. These papers provide an overall picture of a calcalkaline subduction-related magmatic arc which was active from earliest Ordovician (Tremadocian) times. In the south there is only a plutonic record and the arc was clearly developed on the margin of the Gondwana foreland that had been established during the Pampean orogeny (Pankhurst e t ai. this volume). Now, even in the north, the sedimentological evidence discussed by Bahlburg (this volume) appears to favour a continental arc rather than the earlier model of an island arc for the Famatinian environment at this Early Ordovician stage. Subsequently, more evolved and less deformed granitoids, some with a marked peraluminous tendency, were emplaced into the batholithic belt until mid-Ordovician times. Activity then ceased until peraluminous intra-plate magmatism restarted to the east of the old arc, in Devonian times. At this point, the Palaeozoic geological record of the Sierras Pampeanas ceased, indicating the stability of the new continental margin, at least until the start of the Andean cycle in Cretaceous times.

8

R.J.

P A N K H U R S T & C. W. R A P E L A

Interpretations presented in the contributions in this volume have in m a n y cases developed since the 1996 congress in Buenos Aires, and we have not attempted to ensure total consistency between the various contributions. We have simply tried to point out here that radically different views are expressed concerning the d u r a t i o n and extent o f the F a m a t i n i a n arc, the question o f w h e t h e r this arc was an island arc or a continental one, and the timing o f the final collision of the Precordillera terrane with G o n d w a n a . The reader is invited to c o m p a r e these views critically, as the resolution is necessary to reach a fully coherent u n d e r s t a n d i n g of the processes involved in this, one of the best-constrained allochthonous terrane histories k n o w n . The new data referred to in this introduction were obtained as part of contract CII-CT92-0088" from the European Commission and was also supported by grant PIP No 4148 from the Consejo Nacional de Investigaciones Cientificas y Trcnicas de la Repfiblica Argentina (CONICET); this paper is listed as NIGL Publication No. 238. The volume as a whole is a contribution to IGCP Projects 345 (Andean Lithospheric Evolution) and 376 (Laurentia-Gondwana Connections before Pangea).

References ACE/qOLAZA, F. G. & TOSELLI, A. 1973. Consideraciones estratigr~ificas y tectbnicas sobre el Paleozoico inferior del noroeste argentino. 2 ~ Congreso Latinoamericano de Geologia, Actas, 2, 755-763. , MILLER, H. & TOSELLI, A. J. (eds) 1990. El Ciclo Pampeano en el Noroeste Argentino. Universidad Nacional de Tucumfin, Serie Correlaci6n Geol6gica, 4. ASTINI, R. A. 1996. Las fases diastr6ficas del paleozoico medio en la Precordillera del oeste Argent i n a - evidnecias estratigrfificas. XIII Congreso Geoldgico Argentino y III Congreso de Exploracidn e Hidrocarburos, Buenos Aires, Actas, V, 509-526. 1998. Stratigraphical evidence supporting the rifting, drifting and collision of the Laurentian Precordillera Terrane of western Argentina. This volume. --, BENEDETTO, J. L. & VACCARI, N. E. 1995. The early Paleozoic evolution of the Argentine Precordillera as a Laurentian rifted, drifted and collided terrane. Geological Society of America Bulletin, 107, 253-273. BAHLBURG, H. 1998. The geochemistry and provenance of Ordovician turbidites in the Argentine Puna. This volume. & HERVE, F. 1997. Geodynamic evolution and tectonostratigraphic terranes of northwestern Argentina and northern Chile. Geological Society of America Bulletin, 109, 869-884. -

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BENEDETTO, J. L. 1998. Early Palaeozoic brachiopods and associated shelly faunas from western Gondwana: their bearing on the geodynamic history of the pre-Andean margin. This volume. BORDONARO, O., KELLER, M. & LEHNERT, O. 1996. E1 Ordovicico de Pon6n Trehue en la provincia de Mendoza (Argentina): redefiniciones estratigrfificas. XIII Congreso Geolrgico Argentino y III Congreso de Exploracirn e Hidrocarburos, Buenos Aires, Actas, I, 541-550. CAMPOSNETO, M. C. & FIGUEIREDO,M. C. H. 1995 The Rio-Doce orogeny, southeastern Brazil. Journal of South American Earth Sciences, 8, 143-162. CRIADO ROQUI~, P. 1972. Cinturbn mbvil mendocinopampeano. In: LEANZA, A. F. (ed.) Geologia Regional Argentina. Academia Nacional de Ciencias de la Repfiblica Argentina, Cbrdoba, 297-303. DALLA SALDA, L. H., CINGOLANI,C. A. 8,~VARELA,R. 1992a. Early Paleozoic orogenic belt of the Andes in southwestern South America: result of Laurentia-Gondwana collision? Geology, 20, 617-620. - - , DALZIEL,I. W. D., CINGOLANI,C. A. & VARELA, R. 1992b. Did the Taconic Appalachians continue into southern South America? Geology, 20, 1059-1062. --, LOPEZ DE LUCHI, M. G., CINGOLANI, C. A. & VARELA, R. 1998. Laurentia-Gondwana collision: the origin of the Famatinian-Appalachian orogenic belt (a review). This volume. DALZIEL, I. W. D. 1997. Neoproterozoic-Paleozoic geography and tectonics: Review, hypothesis, environmental speculation. Geological Society of America Bulletin, 109, 16-42. , DALLA SALDA, L. n., CINGOLANI, C. A. 8,~ PALMER, P. 1996. The Argentine Precordillera: a Laurentian Terrane? Penrose Conference Report. GSA Today, February 1996, 16-18. DICKERSON, P. W. & KELLER, M. 1998. The Argentine Precordillera: its odyssey from the Laurentian Ouachita margin towards the Sierras Pampeanas of Gondwana. This volume. GRISSOM, G. C., DEBARI, S. M. & SNEE, L. W. 1998. Geology of the Sierra de Fiambalfi, northwest Argentina: implications for Early Palaeozoic Andean tectonics. This volume. GUTII~RREZ-MARCO,J. C., SAAVEDRA,J. & RABANO,I. (eds) 1992. Paleozoico Inferior de IberoamHica. Universidad de Extremadura, Spain. HUFF, W. D., BERGSTROM, S. M., KOLATA, D. R., CINGOLANI, C. & ASTINI, R. A. 1998. Ordovician K-bentonites in the Argentine Precordillera: relations to Gondwana margin evolution. This volume. ISACKS, B. 1988. Uplift of the central Andean plateau and bending of the Bolivian Orocline. Journal Geophysical Research, B93, 3211-3231. JORDAN, T. E. & ALLMENDINGER, R. W. 1986. The Sierras Pampeanas of Argentina: a modern analogue of Rocky Mountain foreland deformation. American Journal of Science, 286, 737-764. KAY, S. M. & GORDILLO, C. E. 1994. Pocho volcanic rocks and the melting of depleted continental lithosphere above a shallowly dipping subduction zone in the central Andes. Contributions to Mineralogy and Petrology, 117, 25-44.

INTRODUCTION

--,

ORELL, S. • ABRUZZI, J. M. 1996. Zircon and whole rock Nd-Pb isotopic evidence for a Grenville age and a Laurentian origin for the basement of the Precordillera in Argentina. Journal of Geology, 104, 637-648. KELLER, M. & DICKERSON, P. W. 1996. The missing continent of Llanoria- was it the Argentine Precordillera? XIII Congreso Geol6gico Argentino y III Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, V, 355-367. - - , BUGGISCH, W. & LEHNERT, O. 1998. The stratigraphical record of the Argentine Precordillera and its plate-tectonic background. This volume. LINARES, E., LLAMBiAS,E. J. & LATORRE, C. O. 1980. Geologia de la provincia de La Pampa, Repfiblica Argentina y geocronologia de sus rocas metam6rficas y eruptivas. Revista de la Asociaci6n Geol6gica Argentina, 35, 87-146. LLAMBiAS, E. J., CAMINOS, R. & RAPELA, C. W. 1984. Las plutonitas y vulcanitas del Ciclo Eruptivo Gondwfinico. In: RAMOS, V. A. (ed.) Geologia y Recursos Minerales de la Provincia de Rio Negro. Relatorio del 9 ~ Congreso Geol6gico Argentino, 83-117. - - - , SATO, A. M., ORTIZ SUAREZ, A. & PROZZI, C. 1998. The granitoids of the Sierra de San Luis. This volume. MCDONOUGH, M., RAMOS, V. A., ISACHSEN, C. & BOWRING, S. 1993. Edades preliminares de circones del basamento de la Sierra de Pie de PaiD, Sierras Pampeanas Occidentales de San Juan: sus implic/mcias para el supercontinente proterozoico de Rodinia. XII Congreso Geol6gico Argentino y II Congreso de Exploraci6n de Hidrocarburos, Mendoza, Actas, III, 340-343. PANKHURST, R. J., RAPELA, C. W., SAAVEDRA, J., BALDO, E., DAHLQUIST, J., PASCUA, I. & FANNING, C. M. 1998. The Famatinian magmatic arc in the central Sierras Pampeanas. This volume. RAMOS, V. A. &BASEI, M. 1997. The basement of Chilenia: an exotic continental terrane to Gondwana during the early Paleozoic. In: BRADSHAW,J. D. & WEAVER, S. D. (eds) Terrane Dynamics-97. International Conference on Terrane Geology, Christchurch, New Zealand, Abstracts, 140-143. , DALLMEYER, R. D. & VUJOVICH, G. I. 1998. Time constraints in the Early Palaeozoic docking of the Precordillera, central Argentina. This volume. , JORDAN, T. E., ALLMENDINGER, R. W., MPODOZIS, C., KAY, S. M., CORTI~S, J. M. & PALMA, M. A. 1986. Paleozoic terranes of the central Argentine-Chilean Andes. Tectonics, 5, 855-880. RAPELA, C. W., PANKHURST, R. J., LLAMBiAS,E. J., LABUDIA, C. & ARTABE, A. 1996. "Gondwana"

9

magrnatism of Patagonia: Inner Cordilleran calcalkaline batholiths and bimodal volcanic province. Third International Symposium on Andean Geodynamics, St.-Malo, France, R~sumbs 6tendus, ORSTOM, Paris, 791-794. - - , CASQUET, C., BALDO, E., SAAVEDRA,J., GALINDO, C. & FANNING, C. M. 1998. The Pampean Orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de Cbrdoba. This volume. RENNE, P. R., ERNESTO, M., PACCA, I. G., GLEN, J. M., PREVOT, M. & PERRIN, M. 1992. The age of Paranfi flood volcanism, rifting of Gondwanaland, and the Jurassic-Cretaceous boundary. Science, 258, 975-979. SAAVEDRA,J., TOSELLI, A., ROSSI, J., PELLITERO, E. & DURAND, F. 1998. The Early Palaeozoic magmatic record of the Famatina System: a review. This volume. SIMS, J. P., IRELAND,T. R., CAMACHO,A., LYONS, P., PIETERS, P. E., SKIRROW, R. G., STUART-SMITH, P. G. & MIR0, R. 1998. U-Pb, Th-Pb and Ar-Ar geochronology from the southern Sierras Pampeanas, Argentina: implications for the Palaeozoic tectonic evolution of the western Gondwana margin. This volume. THOMAS, W. A. & ASTINI, R. A. 1996. The Argentine Precordillera: A traveler from the Ouachita embayment of North America Laurentia. Science, 273, 752-757. TOSDAL, R. M. 1996. The Amazon-Laurentian connection as viewed from the Middle Proterozoic rocks in the central Andes, western Bolivia and northern Chile. Tectonics, 15, 827-842. TURNER, S., REGELOUS, M., KELLEY, S., HAWKESWORTH,C. J. & MANTOVANI,M. 1994. Magmatism and continental break-up in the South Atlantic: high-precision 4~ geochronology. Earth and Planetary Science Letters, 121, 333-348. VARELA, R., LOPEZ DE LUCHI, M., CINGOLANI, C. & DALLA SALDA, L. 1996. Geocronologia de gneises y granitoides de la Sierra de Umango, La Rioja. Implicancias tect6nicas. XIII Congreso Geol6gico Argentino y III Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, III, 519-527. YON GOSEN, W. & PROZZI, C. 1998. Structural evolution of the Sierra de San Luis (Eastern Sierras Pampeanas, Argentina): implications for the proto-Andean margin of Gondwana. This volume. VUJOVlCH, G. I. & KAY, S. M. 1998. A Laurentian? Grenville-age oceanic arc/back-arc terrane in the Sierra de Pie de Palo, Western Sierras Pampeanas, Argentina. This volume.

Stratigraphical evidence supporting the rifting, drifting and collision of the Laurentian Precordillera terrane of western Argentina RICARDO

A. A S T I N I

Estratigrafia y Geologia Hist6rica, Departamento de Geologia, Facultad de Ciencias Exactas Fisicas y Naturales, Universidad Nacional de C6rdoba, Vdlez Sarsfield 299, CC 395, 5000 C6rdoba, Argentina (e-mail: [email protected]) Abstract: Major stratigraphical evidence points toward initial rifting of the Precordillera

terrane of western Argentina during the Early Cambrian from the Ouachita embayment in southeastern Laurentia. A first fully developed rifting episode is suggested from subsidence analysis and the sedimentary record. Restricted graben-fill red-beds and evaporites in the eastern tectofacies, together with coarse arkose and quartz conglomerate preserved as blocks in the western tectofacies, are interpreted as deposits of the initial rift-related brittle extension. Stratigraphical similarities with localities surrounding the Ouachita embayment leads to a model in which the Precordillera was the southern part of the pre-Appalachian Laurentia margin and the conjugate margin of the Texas promontory, separated along the Alabama-Oklahoma transform and the Ouachita rift, respectively. Overstepping of syn-rift facies by a thick succession of Lower Cambrian to Lower Ordovician carbonates indicates completion of rift-to-drift transition and initiation of thermo-tectonic subsidence characteristic of continental terrace passive-margin development. Slope facies, marginal faunal belts and recently dated mafic rocks indicate the development of an ocean basin separating the Precordillera from Laurentia by the Mid-Cambrian. A 20-30 Ma period of drifting and isolation is indicated from faunal and sedimentological evidence. Subsidenceinduced incipient drowning and abundant K-bentonites suggest the influence of subduction and the progressive approach of the terrane to western Gondwana. Continuous Mid- and Late Ordovician post-collisional extension brackets the collision, although no clear evidence of shortening has been detected in the preserved former sedimentary bank. Post-collisional extension is separated by >70 Ma from the initial rifting. The Argentine Precordillera has basement rocks and lowermost Palaeozoic cover (Cambrian and most of the Ordovician) which strongly contrast with those of the surrounding geological provinces in southern South America (Fig. 1). For this reason, it has been considered a suspect terrane and has recently been the target of much interest due to the current controversy on early Palaeozoic palaeogeography and plate tectonics. Although it is accepted that the Precordillera should be considered an allochthonous or exotic terrane (Benedetto 1993; Astini et al. 1995a, 1996; Dalziel et al. 1996), there is much debate on the source and route it took to Gondwana. Two major end-member hypotheses have been suggested to explain the transfer of the Precordillera from Laurentia to Gondwana, each implying different kinematics and timing: (1) the terrane hypothesis largely considers the Precordillera as a rifted-drifted microcontinent (Astini et al. 1995a); and (2) the tectonic tracer model (Dalziel, 1993), where after a continentcontinent collision and later rifting, a sliver of Laurentia was left attached to Gondwana (Dalziel et al. 1994; Dalziel 1997). The contrasts between the two models are much greater than is apparent, with different geological implications

and consequences apart from the origin of the Ocloyic orogeny and the Taconic orogeny in the Mid-Ordovician, respectively in western Gondwana and in eastern Laurentia. Stratigraphical history, including the record of subsidence and tectonic history, and sensitivity to climate change, should be regarded as the critical tool to discriminate between these two models. The original departure of the Precordillera block from Laurentia is undisputed, because its Cambrian faunas (Borrello 1971; Vaccari 1994; Lehnert et al. 1997) and sedimentary successions are nearly identical to those of the continentalshelf successions of Laurentia (Bond et al. 1984; Ramos et al. 1986; Benedetto, 1993; Astini et al. 1995a, 1996). Based on close stratigraphical comparisons and kinematic constraints, Thomas & Astini (1996a) concluded that the Precordillera continental fragment was most probably rifted from the Ouachita embayment of the southern margin of Laurentia. New palaeomagnetic data (Rapalini & Astini 1998) confirm this conclusion. Docking with Gondwana apparently occurred between the Early and Late Ordovician, when the Precordillera stratigraphy (Astini et al. 1995a, 1996), faunal affinities (see Benedetto this volume), and

ASTINI, R. A. 1998. Stratigraphical evidence supporting the rifting, drifting and collision of the Laurentian Precordillera terrane of western Argentina. In: PANKHURST, R. J. & RAPELA, C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 11-33.

12

R . A . ASTINI

Fig. 1. Map of the main Ordovician outcrops of southern South America showing the lithological contrast between the Argentine Precordillera carbonates and the surrounding siliciclastic and volcaniclastic basins (after Astini 1995a).

sedimentary overlap first show a definitive connection with Gondwana. At present, timing for the collision appears to be in the Middle Ordovician, based mainly on palaeo-biogeographical arguments. Benedetto & Astini (1993), Benedetto (1993), Astini et al. (1995a) and Thomas & Astini (1996a) have suggested that between the initial rifting from Laurentia and the accretion to western Gondwana (the present western side of South America) the Precordillera drifted free and isolated in the Iapetus Ocean, where it had some faunal exchange with other intra-Iapetan island arcs and terranes. The purpose of this paper is to summarize the stratigraphical evidence for the allochthonous nature of the Precordillera and the data that have particularly helped in establishing the model of rifting, drifting, and collision of this exotic-to-Gondwana terrane.

Geological setting and previous work The Argentine Precordillera is a Grenville-age lithospheric block (Kay et al. 1996; Astini et al. 1996) with a thick Palaeozoic cover of folded marine to non-marine sedimentary rocks and Tertiary strata (Allmendinger et al. 1990). Post-

Triassic Mesozoic sediments are absent. Structurally, the Precordillera is considered as part of the extensive fold-and-thrust belt located east of the main Andes (Fig. 2). To the east it is bounded by the broken foreland of the Sierras Pampeanas, a metamorphic basement presently considered as late Precambrian-Early Cambrian in age (Kraemer et al. 1995; Rapela et al. this volume; Sims et al. this volume). The Precordillera thrust-fold belt extends for 400 km between Mendoza and La Rioja provinces in western Argentina, but its present-day features are strongly controlled by Andean segmentation (Jordan et al. 1983). Thus, the real north-south length of the exotic block might exceed 800 km if its extension into the San Rafael block, south of Mendoza, is considered (Astini et al. 1995a; Bordonaro et al. 1996; Ramos et al. this volume). The Precordillera terrane is over 200 km wide (Fig. 2) and may include the Grenvillian basement that crops out immediately to the east (see Pankhurst & Rapela this volume) in the Western Sierras Pampeanas. Identification of terranes can be made by considering the regional extent and overlapping relationships of the sedimentary units. The Lower Cambrian to lower Upper Ordovician

THE LAURENTIAN PRECORDILLERA TERRANE

13

Fig. 2. Location map of the Argentine Precordillera in its present location to the east of the High Andes and detail of the eastern and western tectofacies recognised by Astini (1992) for the Ordovician. Note that part of the basement included in the western Pampeanas as well as the outcrops to the south of San Rafael are included in the Precordillera terrane. Localities are numbered. I, 2, 3, 7, 8, 10 and 11 stand for horst-block nucleating carbonates in the Mid- and Late Ordovician.

section, as well as the Grenvillian basement, seem to be restricted to the Precordillera domain, thus constraining the limits of the exotic terrane. Starting in the uppermost Ordovician, sedimentary units and depositional environments (Hirnantian glacial horizon and post-glacial transgressive deposits) can be traced throughout the rest of the Andean and Sub-

andean belt, which gives a minimum age for the proposed accretion of the Precordillera terrane. Different Cambrian-Ordovician tectono-stratigraphical assemblages reflecting major basin evolutionary stages were identified by Astini & Benedetto (1996). A basal, dominantly carbonate, continental platform cycle (including its rifting and drifting history) can be separated

14

R . A . ASTINI

from a complex, mostly siliciclastic, wedge representing the syn- to post-collisional cycle. A two-fold subdivision of the Precordillera into eastern and western tectofacies can be made on the basis of substrate nature, dominant sedimentary composition, and structural behaviour and response (Fig. 2). According to Astini (1992), the eastern tectofacies comprises most of the shallow-marine carbonate stratigraphy and overlying environmentally restricted clastic rocks, whereas the western tectofacies includes a deep-marine siliciclastic assemblage. Important features in the latter are the association of mafic and ultramafic igneous rocks related to remnants of spreading-ridge centres (Kay et al. 1984) and the widespread low-grade metamorphism. The eastern domain substrate may thus be interpreted as normal continental lithosphere (35-40 kin), whereas the western domain is characteristic of thinner crust. The outstanding stratigraphical feature of the Precordillera is the thick Cambrian and Lower Ordovician carbonate succession (Baldis & Bordonaro 1981; Keller et al. 1994; Keller et al. this volume), which contrasts markedly with coeval clastic and volcaniclastic successions in the rest of the South American margin. This, together with the Laurentian trilobite faunas, led Poulsen (1958) and Borrello (1965, 1971) to consider its possible exotic origin (Ramos et al. 1986). More recently, palaeo-biogeographical analysis based on brachiopods led Benedetto (1993; see also Astini et al. 1995a, fig. 7) to suggest that the Precordillera had remained close to Laurentia until the earliest Ordovician. Further detailed taxonomic studies of trilobites allowed Vaccari (1994; see also Palmer et al. 1996) to establish a close comparison with the south-eastern margin of Laurentia (see Astini et al. 1995a, fig. 5). Finally, Astini et al. (1995b), on stratigraphical grounds, pointed out the very close match between the northern Precordillera and the Alabama Appalachians. Astini (1995a) elaborated a wander path based on lithologically sensitive climate indicators, showing that the Precordillera remained in low latitudes ('Pacific realm') during the Cambrian and earliest Ordovician and then migrated toward temperate environments, with clear indications of the Gondwana glaciation by Late Ordovician (Hirnantian) time. Thus, while the parent continent Laurentia remained within the tropics, the Precordillera terrane drifted southward, and by the early Late Ordovician it occupied a relatively high southern latitude. This suggests that the Precordillera terrane was most probably isolated from Laurentia long before the onset of the Ocloyic Orogeny.

In the past there has been little palaeomagnetic work in the Lower Palaeozoic rocks of the Argentine Precordillera. Now, for the first time, primary magnetisation has been isolated by Rapalini & Astini (1998), at several sites in northern Precordillera. This study confirms the suggestion that the Precordillera constitutes an exotic-to-Gondwana terrane, and its position during the Early Cambrian is remarkably coincident with that suggested by Astini et al. (1995a) and Thomas & Astini (1996a).

Hypotheses and background Because the original rifted margins of both the Appalachian-Ouachita region and the Precordillera have been deformed by later orogenies and at present are partly covered by younger rocks, details of the rift site and geometry, as well as the paired rifted-margin palaeogeography, are difficult to determine. For this reason, alternative interpretations have been suggested for the relative timing of separation from Laurentia and of accretion to Gondwana, as well as of the exact positioning and kinematics of rifting (e.g. Dalla Salda et al. 1992a; Dalziel et al. 1994; Dalziel, 1997; Thomas & Astini 1996a; Astini et al. 1996; Keller & Dickerson 1996; Buggisch 1996). Current discussion focuses on whether or not the Precordillera represented a 'bridge' or a 'raft' between the two principal Early Palaeozoic land masses. This, in turn, is important to deciding between the conventional (archetypal) and alternative plate palaeo-geographical schemes (see Torsvik et al. 1995), in view of the longitude uncertainty in palaeomagnetically based reconstructions. Much evidence has been put forward to support isolation of the Precordillera in an intra-Iapetus position before it collided with Gondwana (Astini et al. 1995a, 1996; Thomas & Astini 1996a; Dalziel et al. 1996). However, in seeking a longitude constraint for the wandering of Laurentia in the 'end-run' hypothesis (Dalziel 1991), an alternative solution to the traditional northern hemisphere continental puzzle was suggested by Dalziel (1997) as a modification of his earlier work (Dalziel et al. 1994). A keystone in this most recent reconstruction is the Argentine Precordillera, which is proposed to have had limited separation from Laurentia through initial crustal extension in the Early Cambrian. In this model the Precordillera block would have had to move outboard of the Ouachita embayment to occupy an 'intraIapetus' position (but still intra-shelf), analogous to the present-day Maurice Ewing bank

THE LAURENTIAN PRECORDILLERA TERRANE southwest of the Falkland/Malvinas plateau in the South Atlantic. According to Dalziel (1997) this model explains the increasing isolation of the Precordillera, while still united to Laurentia through a continental bridge (extended Grenvillian crust); the Precordillera 'peninsula' would have collided with Gondwana. Only after collision did Dalziel suggest development of oceanic crust within the western Precordillera clastic wedge, interpreted as the result of rifting between Laurentia and Gondwana. The main difference from the microcontinent model of Astini et al. (1995a) and Thomas & Astini (1996a) is that (as previously suggested in Dalla Salda et al. 1992a, b) the Precordillera would have drifted as an attached part of Laurentia. Hence, both would have shared the same polar wander path during the Early Ordovician, approaching that of Gondwana. In contrast, Astini et al. (1995a) and Thomas & Astini (1996a) considered the Precordillera block as an independent microcontinent, so that the Laurentian and Gondwanan paths would have never converged to indicate a common Ocloyic-Taconic orogeny. During supercontinent break-up, smaller continental fragments surrounded by various rift branches and transfer faults (later transforms) are left between the larger continents (e.g., Rosendahl 1987; Etheridge et al. 1989). A rifted microcontinent may remain near the parent continent and, during a subsequent collision, be reattached to the parent continent (e.g., the Appalachian internal basement massifs of Thomas 1977 and Hatcher 1983). Alternatively, the rifted lithosphere fragment may separate fully from the parent continent and drift independently to collide with a foreign continent (e.g., Howell 1989). In this case it would be interpreted as a far-travelled terrane. Rift-related intracratonic fault systems indicate brittle crustal extension of the parent continent (Thomas 1991). Failed rifts or aulacogens partially or completely outline basement blocks (Braile et al. 1986), indicating interrupted microcontinental evolution due to incomplete break-up of the parent continent (e.g., SE Laurentia). In the context of rifted margins, a fragment of Grenvillian continental crust, now recognised as the Precordillera, appears to have rifted from the Ouachita embayment of the Laurentian margin during the Early Cambrian (Thomas 1991), as a result of major diachronous rifting associated with the opening of the Iapetus. Three major lines of stratigraphical evidence support the microcontinental model for the Precordillera: (a) the stratigraphical record of rifting and passive margin history, (b) the

15

stratigraphical record of drowning and progressive flexure recording foredeep and forebulge formation in a peripheral foreland during accretion, and (c) the stratigraphical record of major palaeogeographical rearrangement related to widespread extension after collision. Faunal evidence provides precise constraints on the timing and degree of isolation (Benedetto this volume).

Stratigraphical record of rifting and drifting E a r l y C a m b r i a n - E a r l y Ordovician carbonate bank

As previously mentioned the striking feature of the Argentine Precordillera is the Early Cambrian-Early Ordovician marine carbonate succession, which is nearly 3000m thick. This was originally recognized as a typical passive margin peri-cratonic succession by Gonzfilez Bonorino (1975) and Baldis & Bordonaro (1981, 1985), and consists of several units, mainly of shallow to very shallow marine origin (Fig. 3) (Keller et al. 1994). Lower to Middle Cambrian thick, shoaling-upward, carbonate units, alternating with siliciclastic rocks (cf. Aitken 1966; grand cycles according to Baldis & Bordonaro 1982) are succeeded by thick dolostones with sparse stromatolite-bearing horizons. These, in turn, grade into a rhythmically arranged thick- to medium-bedded, mostly dolomitic Upper Cambrian unit composed of recurrent small-scale, shallowing-upward peri-tidal cycles (Keller et al. 1989; Cafias 1990; Armella 1994) and sub-tidal cycles (Cafias 1995a). An exposure-and-drowning surface separates this from the lowermost Ordovician unit, which, according to Cafias (1995b) corresponds to a typical Bahamian restricted platform, bordered by grainstone shoals toward the open sea. The uppermost unit in the carbonate bank is a calcareous, dominantly sub-tidal, muddy succession with a rich open-marine calcareous fauna, punctuated by sea-level fluctuations responsible for the generation of reef-mounds and biostromes (Cafias & Carrera 1993; Cafias & Keller 1993; Keller & Bordonaro 1993). Similar facies types of the Cambro-Ordovician carbonate bank and vertical arrangement compared with the coeval Laurentian carbonate rim, led Astini et al. (1995a) to favour correlation of the Precordillera with the southern Appalachian platform. By the Late Cambrian, the Precordillera had developed into an agraded, flat-topped rimmed-platform dominated by peritidal facies (Cafias 1995a; Armella 1994; Astini et al. 1995a), similar to that interpreted for the

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Fig. 3. Generalized time-stratigraphic chart of the Cambrian and Ordovician across the Argentine Precordillera. Note contrasting stratigraphy of the eastern and western tectofacies (see distribution in Fig. 2). CTF, Cerro Totora Formation; LLF, La Laja Formation; ZF, Zonda Formation; LFF, La Flecha Formation; LSiF, La Silla Formation; SJF, San Juan Formation; GF, Gualcamayo Formation; ChF, Las Chacritas Formation; LAF, Las Aguaditas Foramtion; LPF, Las Plantas Formation; LVF, Las Vacas Formation (La Cantera Formation); SF, Sassito Formation; TF, Trapiche Formation; DBF, Don Braulio Formation; LChF, La Chilca Formation; LSF, Los Sombreros Formation; YLF, Yerba Loca Formation; AF, Alcaparrosa Formation. Time-scales according to Tucker & McKerrow (1995). Olistoliths and basement boulders (granites and metamorphic rocks) present in the Los Sombreros Formation (western tectofacies) are represented according to their individual ages (Benedetto & Vaccari 1992; Bordonaro & Banchig 1996; Astini unpublished). (Sources for construction: Albanesi et al. 1990; Astini 1991; Astini & Brussa 1997; Astini & Cafias 1995; Astini & Vaccari 1996; Benedetto & Herrera 1987; Brussa 1994; Keller et al. 1994; Lehnert 1995; Ortega & Brussa 1990; Ortega et al. 1991; Vaccari 1994). central-southern Appalachian passive margin (Demicco 1985; Koerschner & Read 1989; Osleger 1993). The remarkable similarities indicate analogous substrates and latitudes to those of the southern Appalachians. Studies of internal architecture and lateral correlation between Precordilleran units are still limited, although neither very proximal nor really distal environments are well represented in the Argentine Precordillera. It is reasonable then to suggest that the Precordillera is somewhat incomplete, having been a platform at least 400-600km wide, when compared with the similar Appalachian setting (Cafias 1995a; Astini e t al. 1996; Keller 1996). Accordingly,

the assumption that the Precordillera terrane may have been c. 800 km wide (Thomas & Astini 1996a) seems to be reasonable. E - W shortening may have greatly affected the Precordillera terrane in post-Early Ordovician time, stripping the carbonate cover from the Grenvillian basement that now crops out to the east of the Precordillera thrust belt. Tectonic reworking also has affected and removed clear evidence of the western talus (present co-ordinates), which has been greatly shortened and overprinted by both Andean and Early Devonian tectonism (Astini 1996). Nevertheless, an E - W passive margin polarity is the consensus among the Precordilleran workers, based on most of

THE LAURENTIAN PRECORDILLERA TERRANE

17

the olistoliths and slumped blocks of the western tectofacies (Fig. 3) being interpreted as former slope and platform edge facies, gravitationally reworked into more distal environments. Using data from Baldis & Bordonaro (1981), Bond et al. (1984) compared the Precordillera subsidence history with that of the various Laurentian margin domains and concluded that it had most similarity with that of the Appalachian platforms. Therefore, Bond et al. (1984) considered the Precordillera as a probable Appalachian conjugate margin. According to their continental break-up hypothesis, the Precordillera was first positioned as a centralnorthern Appalachian counterpart, a scheme later adopted by Herrera & Benedetto (1991). Nevertheless, Bond et al. (1984) also suggested the alternative that the Precordillera be considered a displaced terrane that was part of the Appalachian miogeocline in the earliest Palaeozoic, a scheme favoured by Ramos et al. (1986) and Benedetto (1993). E v i d e n c e f o r rifting

One premise of similar thick carbonate platforms, such as that of the Precordillera, is that they have always been recognized as typifying low-latitude stable passive-margin continental terrace deposits (e.g., Bond et al. 1983, 1995). These carbonate edifices reflect the thermal subsidence of a continental margin after rifting and hence indirectly indicate the age of break-up of the margin (Bond et al. 1984). The presence of such a typical passive margin carbonate bank implies a minimum thickness of continental crust, sufficient to maintain shallow bathymetry. The first-order control on subsidence and growth of the Precordilleran carbonate platform, as well as for the Laurentian analogues, was thermal contraction of heated lithosphere following rifting (Bond et al. 1989). This argument can be taken as a powerful indicator of fully developed Early Cambrian rifting between the Precordillera and the Ouachita margin. Recent work in the northern Precordillera (Astini & Vaccari 1996) identified a new Lower Cambrian unit, the Cerro Totora Formation (Fig. 3), composed of a mixed succession of evaporites, red siliciclastic rocks and carbonates (Fig. 4), underlying the thick Cambrian to Lower Ordovician carbonates (>2500m). The Cerro Totora Formation records the transition from syn-rift evaporites and marginal-marine red or variegated arkosic sedimentary rocks to quartz arenites and intercalated olive green/ glauconitic shales that reflect generalised transgression associated with the cessation of rifting.

Fig. 4. Stratigraphic section of the syn-rift clastic and evaporitic graben fill succeeded by limestones and dolostones after the main transgressive surface that represents the rift-drift transition in the northern Precordillera (modified from Astini & Vaccari 1996). The basal anhydrite-gypsum succession (>250m) is a common depositional facies in the early stages of restricted-circulation divergent margins in low-latitude arid-climate settings (Kinsman 1975; Rona 1982; Warren 1989). The evaporites are strongly recrystallized and interbedded with dolomitized crypto-microbial and oolitic tabular limestones, and are succeeded by a predominantly siliciclastic interval (c. 50m). This red-bed interval is composed of

18

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red to purplish silty shales and minor quartzose to sub-feldspathic sandstones and intra-clastic conglomerates. Thin beds of reworked detrital evaporites and carbonate layers also occur. Halite hopper-cube casts and intra-sedimentary gypsum crystals indicate extensive evaporation in arid supra-tidal environments, as shown by modern analogues (Shin 1983; Handford 1991). Mud-cracked horizons, tepee structures and brecciated horizons are also common features, suggesting subaerially exposed mud-flat and salt-flat environments (Kendall & Warren 1987). Facies analysis, together with the general stratigraphical development, allows interpretation of this unit as a typical syn-rift succession (Astini & Vaccari 1996). Starting in the late Early Cambrian, the onset of the passive-margin stage is recorded in the upward transition to the thick carbonate bank. Astini et al. (1995b) pointed out the time parallelism and the nearly identical lithology of

the Cerro Totora Formation with the contemporaneous stratigraphy of the Rome (Waynesboro) Formation and related units in the southern Appalachians and Black Warrior Basin, where the rifting history developed similar graben-filled successions (e.g., Mississippi Valley trough and Birmingham graben system, Mellen 1977; Thomas 1991; Raymond 1991). On this basis, Thomas & Astini (1996a) concluded that the Precordillera block was rifted from the Ouachita embayment, initially bounded by the Blue Ridge rift to the east, the Alabama-Oklahoma transform to the north and the Ouachita rift to the west (Fig. 5). NWpropagation of the Alabama-Oklahoma transform fault apparently transferred extension from the mid-Iapetus ridge, along the Blue Ridge rift to the northern part of the Ouachita rift, initiating effective separation of the Precordillera block (Fig. 5). The Alabama-Oklahoma transform fault and the Ouachita rift, outlining the

/

r

,,4'

1

Rifled margin of continental crust Transform fault Intracratonic basement fault ,, ,, Cratonward limit of AppalachianOuachita detachment ,, Thrust fault . . . . . . . Margin of Gulf and Atlantic Coastal Plains

Atlantic Ocean

!\ Gulf of Mexico

•N I

Scale 200 km

~-~ . . a

Fig. 5. Plan view of the southeastern USA with locations of the Ouachita embayment of Laurentia, the Argentine Precordillera block and the general outline of the Late Precambrian-Cambrian rifted margin and related structures (after Thomas & Astini 1996a).

THE LAURENTIAN PRECORDILLERA TERRANE Ouachita embayment and the Texas promontory on the Laurentian margin, are defined by deep borehole and geophysical data (summarized in Thomas 1991). Syn-rift igneous rocks range between 552-4-7 Ma and 525 + 25 Ma. Two intracratonic graben systems northeast of the Alabama-Oklahoma transform (the Mississippi Valley and the Birmingham graben systems) indicate extension of continental crust parallel to the transform fault. The Cerro Totora graben (Astini et al. 1995b; Astini & Vaccari 1996) in the northern Precordillera could be considered an analogue to those developed in the Laurentian plate. In the Precordillera, as well as in the Laurentian examples (Thomas 1991), post-rift carbonates seem to overlap the graben-filled successions and the graben-boundary faults, suggesting the end of the brittle stretching phase of rifting and the onset of drifting dominated by thermo-tectonic subsidence. A fundamental line of evidence in continental margins is the recognition of the riftdrift transition (Scrutton 1982). This corresponds to a time of major change in basin tectonics, from block-faulting and graben formation during extension, to regional subsidence controlled by lithospheric cooling and sediment loading at the onset of sea-floor spreading (Bond et al. 1995). Although identification of the break-up unconformity is not always straightforward, in the Precordillera terrane as well as in several sites of the southeastern United States, the transgressive contact (Fig. 4) may represent the rift-drift transition. This transition is late Early Cambrian in the Precordillera, the age of the first clear transgressive limestones that yields open-marine inner-shelf Olenellid trilobites (Astini & Vaccari 1996). Furthermore, rift-drift completion can be deduced from the overstep of continental and transitional syn-rift facies by Mid- and Late Cambrian passive-margin facies (Thomas 1991; Astini & Vaccari 1996). This is a very strong argument in favour of rifting and sea-floor spreading associated with the Ouachita rift, which supports the hypothesis of complete rifting and progressive separation of the Precordillera during the Cambrian. New support for Early Cambrian rifting comes from recent observations on the mafic and ultramafic rocks in the western tectofacies, formerly considered to be 'Famatinian ophiolites' (Haller& Ramos 1984) of Late Ordovician age. The preliminary crystallization age of a mafic ophiolite section in the south-west Precordillera is 565 + 45 Ma (Davis et al. 1997) and indicates, according to these authors, early stages of ocean spreading in the latest Proterozoic, as the Precordillera began to rift from Laurentia.

19

Subsidence analysis, when used as the basis of palaeogeographical reconstructions (e.g., Bond et al. 1984), does not usually consider different rifting models, wherein the thermal history varies greatly. According to Thomas & Astini (1996b, in press), the stratigraphical and kinematic constraints of rifting favour an asymmetric rifting model for the separation of the Precordillera and the Texas promontory. The asymmetric simple-shear rifting model predicts that lower-plate margins will develop thicker carbonate banks, whereas the carbonate fabric on the upper plate would be delayed due to thermal doming. For this reason opposite margins of a conjugate rift pair may not necessarily represent a mirror image as was suggested by Bond et al. (1984). Moreover, asymmetry is one of the most convincing features to support simple-shear detachment along rifted margins (e.g., Etheridge et al. 1989). The detachment model (Lister et al. 1991) predicts that when ongoing extension leads to continental break-up and the formation of an ocean basin, the resulting opposing conjugate passive margins display contrasting structural and uplift-subsidence histories. A passive margin was established around the Ouachita embayment by latest Mid- to Late Cambrian time (Thomas 1991). Thermal doming beneath the Ouachita margin prevented development of syn-rift rocks and notably retarded passive-margin sedimentation. Uppermost Middle Cambrian strata unconformably overlie basement rocks, indicating broad uplift of the margin during rifting (passive-margin mountains), probably related to thermal buoyancy. Local palaeo-topographical relief on the sub-Cambrian unconformity is >200m (Barnes et al. 1972). The Upper Cambrian-Lower Ordovician passive margin succession is c. 1000m thick across the Texas promontory, indicating minor subsidence of the crustal column in post-rift stages. In contrast, a >2500m thick carbonate bank overlying locally important graben-fills developed from late Early to early Mid-Cambrian time. The combination of a narrow zone of transitional crust, a steep slope at the margin of continental crust, the lack of synrift rocks, passive-margin strata directly overlapping the basement, and lower rates of post-rift subsidence, suggests an upper-plate configuration for the Ouachita rift (Thomas 1993). In contrast, the combination of a thick passive-margin succession indicating higher rates of post-rift subsidence and a longer subsidence history, overstepping of graben fill towards the continental block interior, and apparent early stage syn-rift rocks of an extended crust in the

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western tectofacies (quartz-rich conglomerate blocks) suggests that the Precordillera behaved as a lower-plate margin (Thomas & Astini 1996b, in press). Moreover, a lower-plate configuration for the Precordillera terrane would likely explain the diachronous stratigraphical development on the two sides of the rift and the strong stratigraphical similarities between the Precordillera and the Southern Appalachians. The absence of igneous activity on both could reflect their inherited continuity. A major carbonate rim surrounded much of North America during Cambrian and Early Ordovician times, due to the general stability of the craton and its average equatorial position (Sloss 1963; Witzke 1990), although the subsidence history and thermo-tectonic evolution along the margin differ depending on the inherited margin geometry and the kinematics of the different extended crustal blocks during and after rifting (e.g., Thomas 1993). The Precordillera block had its own evolution in terms of subsidence associated with a lowerplate configuration and recorded the rift-drift transition in the late-Early Cambrian. After the Late Cambrian, the Precordillera block apparently moved beyond the Alabama promontory, breaking all connections with Laurentia. Subsequent drift served to isolate the microcontinent. Because of its low-latitude location, a warm climate (Pacific realm) persisted through the earliest Ordovician, producing reef communities and allowing growth of warm water inhabitants.

Stratigraphical signals o f isolation during the Early Ordovician

Continuous exponential-decay thermal subsidence inferred from subsidence analysis (Bond et al. 1984) indicates motion away from the heated rift zone, leading to progressive separation of the Precordillera terrane. This explains the progressive faunal isolation noted by Benedetto (1993) and is strongly supported by palaeo-biogeographical analysis (see Benedetto this volume). The Early Ordovician limestones, in contrast to those of the Late Cambrian, contain no detrital components (Astini et al. 1995a; Cafias 1995a), although neither is there any evidence to suggest an extra-basin source for siliciclastic rocks of the Cambrian units. This may be interpreted as indicating major sea-level fluctuations or the existence of a barrier between the Precordillera bank and continental Laurentia. Scarce detrital content in similar Appalachian

settings is explained by extremely flat topography and wide platform formation with intrabasin depocentres (cf. Markello & Read 1982; Demicco 1985). Although possible, the simplest explanation is to consider the Early Ordovician Precordillera carbonate platform as completely isolated from the continent and lacking relief, similar to the present-day Bahamas (Cafias 1995a, b; Astini et al. 1995a, 1996). Even with sea-level fluctuations as large as those apparently affecting the Precordilleran bank (Keller et al. 1994; Cafias 1995a), continental clastic material would have never reached the Precordillera because of its isolation. Palaeontological criteria (see Benedetto this volume), in addition to those suggested by Astini et al. (1995a, 1996), Benedetto et al. (1995), Vaccari (1995) and Thomas & Astini (1996a), are provided by Lehnert et al. (1997) and Albanesi & Barnes (1997), who remark that the evolution of conodont faunas suggests that the Precordillera was isolated for some time in the Early Ordovician so as to generate endemism, allopatric speciation, and changeover of faunas.

Stratigraphicai record of drowning and progressive flexuring Drowning of the Precordilleran carbonate bank during early Mid-Ordovician time is recorded by progressive westward and southward onlapping of black shales onto the carbonates (Fig. 3) (Astini 1995b; Astini et al. 1995a). Diachronous drowning indicates an increase in water depth across the platform, caused by lithospheric flexure as the Precordillera approached the east-dipping subduction zone beneath the western margin of Gondwana. Variations in thickness of the black shales along strike and across the platform might be related to brittle flexuring, and consequent block-faulting (e.g., Bradley & Kidd 1991). Forebulging towards the western border of the eastern tectofacies (Fig. 2) hindered black shale deposition. There (e.g., Cord6n de Las Chacritas, Mogotes Azules) a small break in sedimentation and keep-up carbonate successions with dramatic shallowing (grainstones) indicate the swelling topography. Progressive intercalation of ash beds, now K-bentonites, in the mid Arenig-early Llanvirn interval (see Huff et al. this volume) serves as independent evidence of arc-related volcanism, most probably related to progressive proximity of the Precordillera to the overriding hanging wall of the orogen, which could have been either an island-arc complex or an active volcanic-arc

THE LAURENTIAN PRECORDILLERA TERRANE seated on western Gondwana margin (the Famatina System?). Although Keller & Dickerson (1996) have suggested their relation with probable bentonites in the Marathon area, considering them as of intraplate extensional origin, there is clear evidence now that there is no counterpart of the Precordilleran K-bentonites in North America, which in nature and timing (mostly Caradoc, Mohawkian of the American subdivision) match more closely those of northwest Europe. The provenance of coarse siliciclastic conglomerates and sandstones (e.g., the Las Vacas and La Cantera formations) (Fig. 3) suggests a source in easterly uplifted areas, most probably related to the Ocloyic orogeny and closure of the marginal ocean basin between the Precordillera and Gondwana. Little is known about direct evidence of shortening linked to the collision of Precordillera. Recent work by Vujovich (1995), and Ramos et al. (1996) recognizes important mylonitic belts affecting the Grenville basement on the eastern border of the Precordillera terrane, close to the boundary with Ramos' (1995) Pampia terrane. Subsurface analyses by Cominguez & Ramos (1991) and Zapata (1996) have also suggested typical collisional zone mylonite fabrics affecting the Grenvillian basement to the east of the Precordillera. Ramos et al. (1996) provided radiometric dates (c. 464Ma) on an associated peak metamorphism during the MidOrdovician (Llanvirn). Associated magmatism in the Sierras Pampeanas yields Early to Late Ordovician ages (c. 490-455 Ma) (see Pankhurst et al. this volume). Finally, west-directed thrusting affecting the passive margin limestones is known in the eastern Precordillera, and has been ascribed to pre-Carboniferous deformation (von Gosen 1992), most probably Ocloyic.

Stratigraphical record of post-collisional rearrangement Astini (1991, 1995c) noticed an important palaeogeographical rearrangement of depocentres and source areas throughout the Mid- to Late Ordovician (Fig. 3). Most of the units in this interval were included in the Las Plantas and the Trapiche alloformations that represent unconformity-bounded heterogeneous mosaics of varied-calibre siliciclastic and carbonate units. Limited lateral continuity, abrupt facies changes (strong lateral grain-size and compositional variations), development of local hiatuses, and restricted sedimentation are some of the common characteristics of this stratigraphical interval. Extensional tectonics have long been

21

used to explain the complex architecture of the Mid- to Late Ordovician Precordilleran stratigraphy (see Borrello 1969). Three different kinds of end-member deposits can be differentiated in the Mid- to Late Ordovician rocks of the Precordillera: (a) coarse- and medium-grained siliciclastic deposits, mainly of extra-basin origin, (b) fine-grained, mostly distal and/or pelagic strata, signifying restricted circulation and slow settling in isolated depocentres, and (c) carbonate intra-basin sedimentation. Although the clastic deposits of both eastern and western tectofacies are mostly extrabasin, and have probably been funnelled to the basin through transfer zones or tectonic troughs, there is a minor basement component derived from intra-basin exposure and erosion of local basement highs. Such is the case for granite boulders in the basal debris flows of the Los Sombreros Formation, the migmatite and schist blocks and pebbles mixed with abundant coarse limestone debris in the La Invernada Formation in the western Precordillera, and the granite and gneiss cobbles and pebbles mixed in the basal conglomerate in the Pon6n Trehue high. The Pon6n Trehue region in southern Mendoza, >200km south of the typical outcrops of the Precordilleran thrust-fold belt (Fig. 2), was formally incorporated into the Precordillera terrane by Astini et al. (1995a), although Baldis & Blasco (1973), Heredia (1982) and Baldis et al. (1985) had previously noted that this region had strong affinities with the Precordillera. Recently, Bordonaro et al. (1996) reinterpreted the local stratigraphy, defining different units in the area that unconformably overlie Grenvillian basement (see Astini et al. 1996). Although the lower unit defined by these authors (Tremadoc-Arenig in age, Lehnert et al. in press) is treated as inplace, following Heredia (1996) I suggest that these limestone outcrops are fallen blocks (olistoliths), which are very common in the middle Upper Ordovician interval (e.g., San Isidro, Los Tflneles, Fig. 2). These limestones are embedded in conglomeratic matrix with abundant basement source components, and the basal conglomerates in nearby outcrops of demonstrated upper Llanvirn-lower Caradoc age have similar composition. Unfortunately, due to later tectonism, it is difficult to trace either outcrop laterally, although this might be ascribed to abrupt depositional topography, with horsts and graben. Also, where exposed, the contacts of the olistoliths with the basement rocks usually display strong shearing. In considering these blocks as in-place, Keller (1996) assigned them to a cratonic section due to their reduced thickness compared to the

22

R . A . ASTINI

expanded sections to the north (the typical Precordilleran outcrops). Nevertheless, no clastic content has been detected in the blocks: hence, it is more logical to consider them as samples of a 'condensed' section in a less subsiding part of the basin. They indicate that even during the Ordovician there was some N-S (present co-ordinates) variation in the Precordilleran carbonate bank architecture. The Pon6n Trehue limestones (Baldis & Blasco 1973) have been correlated on the basis of faunal content and similar lithologies with the Las Aguaditas Formation, south-west of Jfichal (Fig. 2), one of the areas considered as isolated carbonate remnants nucleated on horst blocks. The sedimentary succession of the Pon6n Trehue block reflects a deepening-upwards trend with greywackes and silty shales overlying mixed composition siliciclastic-carbonate conglomerates, bioclastic grainstones and deep-water dark carbonates. Several other carbonate remnants are detected along the central Precordillera (Fig. 2). These include the San Isidro block, the Sassito block, the Los Blanquitos (Las Aguaditas), the Potrerillo Block, the Las Plantas block;

Middle Silurian stata Sassito Formation [late Caradocearly Ashgill] San Juan Formation E (Arenig] o

Jnsgressive conglomerate erosive unconformity

.4,,,,, forced regression [sudden shallowing] ransgressive conglomerate ~""" [sudden deepening]

I

~. ~

skeletal wackestones ~

I[~l with stromatoporoids intrabasinal ~

carbonate

they can be seen as separated blocks or as more or less aligned horsts. Most reveal some kind of active carbonate fabric, nucleated on bottom swells that reached the photic zone under relatively low detrital inflow. This is characteristic of isolated horsts surrounded by relatively deep-water troughs (grabens) that functioned as traps for clastic material. The Pon6n Trehue block is perhaps one of the most remarkable examples, where both basement and overlying strata have been exposed. Two main sedimentary features, indirectly related to block-faulting, are drowning unconformities and forced regressions (Astini 1997a). Drowning unconformities (Schlager 1989) are relatively abrupt terminations of carbonate fabric induced by relative sea-level rise. Conversely, forced regressions (Posamentier et al. 1992) typically indicate rapid sea-level drops. Both features are developed in several localities during the Mid- to Late Ordovician interval. Similar tectonic drowning unconformities are very commonly observed in the Mesozoic Tethyan platform (Bernoulli & Jenkyns 1974; Santantonio 1994) and are also related to widespread

conglomerate

//

erosive unconformity

calcarenites with ~ minor clastic content black shales

~

green siltstones and sandstones chert & phosphate conglomerate

Fig. 6. Sketch showing the stratigraphic relationships and trends observed in the Sassito Formation type section, cropping out in the Sassito creek along the San Juan river (modified from Astini & Cafias 1995). 1, Exposure and erosion (locally developed conglomerate); 2, rapid submergence and restriction (relative sea-level rise); 3, sudden regression and development of shoreface facies (relative sea-level drop); 4, exposure and erosion; 5, renewed transgression and submergence.

THE LAURENTIAN PRECORDILLERA TERRANE extension. Abrupt downward movement of faulted blocks usually generates incipient or full drowning (e.g., the La Chilca block), whereas upward movement of a relatively deep substrate generates abrupt shallowing and, occasionally, produces exposure and erosion (e.g., the Pon6n Trehue block). Some blocks in the Precordillera even reveal a 'yo-yo' tectonic record (cf. BoseUini 1989), having been alternately exposed to erosion and submerged several times. Such is the case of the Sassito block (Fig. 6), which records deep erosion into Arenig limestones, sudden drowning and smothering by black-shale deposition and abruptly overlying shallow shoreface ramp facies (Astini & Cafias 1995), indicating a forced regression. The development of isolated carbonate remnants is perhaps the most straightforward evidence suggesting extension (Fig. 7). Elevated and relatively isolated blocks have served as ideal places for the nucleation of carbonate

23

sediments that might keep up, catch up or drop out according to relative sea level. Fallen blocks, slumped features, intra-formational scars and debris flows are convincing evidence of escarpments and high-gradient slopes surrounding these highs. Hemipelagic carbonates (e.g., the Las Aguaditas Formation) are also common in this environment. Conglomerate and sandstone fills are widely developed in the easternmost grabens in the Precordillera. In south-west Guandacol (northern Precordillera, Fig. 2), angular unconformities and wedging geometry allow reconstruction of half-graben geometry (Fig. 8). Fine-grained black shales in turn indicate restricted sediment inflow from surrounding highs or extra-basin funnelling and stratification of the water column. These are common attributes of isolated grabens (Fig. 7). Maximum stretching, and consequent block faulting, were apparently achieved in the early

Fig. 7. Environmental reconstruction for the main carbonate and restricted facies associations related to horsts and grabens that characterize the Mid- and Late Ordovician extension in the Precordillera (modified from Santantonio 1994). (a) Half-graben development. (b) Horsts and graben development with isolated restricted depocentres in between.

24

R . A . ASTINI

local source escarpment boulders (intrabasinal)

longitudinal or transversal source through canyons or transfer zones (extrabasinal) J stratal wedging ,.

/~ ,k, master normal fault active during sedimenation

hanging-wall e l

1-5 km I

local source (intrabasinal)

rotated block (foot-wall)

I

Fig. 8. Wedged coarse siliciclastic rocks and fallen olistoliths filling an Upper Ordovician hemi-graben in the region to the south-west of Guandacol (northern Precordillera). Note internal angular unconformities and different sources of the infill (after Astini in press). Late Ordovician (Caradoc) when thick massive wedged conglomerates of extra-basin origin were deposited in the eastern tectofacies, and deep-seated extension led to the generation of important volumes of basaltic pillow-lavas in the western Precordillera (Fig. 9). The origin of these mafic submarine volcanic rocks, interlayered with mid-Caradocian graptolitic shales (Blasco & Ramos 1976; Ortega et al. 1991), has been a matter of debate (Kay et al. 1984; Hailer & Ramos 1984, 1993; Ramos et al. 1986; von Gosen 1992; Dalla Salda et al. 1992a; Astini et al. 1996). Based on stratigraphical evidence it seems most reasonable to consider the MORBlike mafic rocks in the Alcaparrosa Formation and associated units of the western tectofacies as products of submarine volcanism related to some kind of incipient spreading-ridge axis,

distinct from the older and real oceanic lithosphere exposed to the south-west of Precordillera (Davis et al. 1997). Recently, Astini et al. (1995a) and von Gosen et al. (1995) have presented explanations for widespread extension, although they argue different contexts for its occurrence. Whereas Astini et al. consider it to be post-collisional, von Gosen et al. visualize it as pre-collisional, related to passive margin extension. Due to its separation from the Cambrian rifting stage (more than 70 Ma apart) and because a thick stable carbonate platform succession separates both stages, it is difficult to interpret this extension episode as related to episodic stretching of the passive margin. Instead, it is proposed that this extension has a different origin related to post-collisional relaxation (Astini 1997b).

Fig. 9. Schematic east-west cross section showing the block-faulted setting for the Precordillera during the Midand Late Ordovician, herein suggested as the main control on sedimentation and architecture for that time interval.

THE LAURENTIAN PRECORDILLERA TERRANE Because syn-collisional extension seems to be rather exceptional, is better interpreted in terms of post-docking disruption (cf. Howell 1989). A major change in plate motion after docking could have been responsible for triggering extension, particularly near the edge of continental lithosphere, hence producing masswasting deposits such as those observed in the western Precordillera (Astini 1991; Keller 1995) and pillow lavas associated with later deepseated extension. In contrast, in the central Precordillera (eastern tectofacies), there seems to be some overlap of both syn-collisional brittle flexural subsidence and post-collisional relaxation. Elastic flexural subsidence would better relate to the relative proximity of Precordillera and Gondwana during the peripheral foreland stage as previously suggested. This would have happened between the mid-Arenig and the early Llanvirn, whereas post-collisional relaxation would correspond to the more severe blockfaulting that occurred during the late LlanvirnCaradoc. Fault overstepping cannot be observed in outcrop so as to infer the end of this extensional episode, but the Late Ordovician units (e.g., the Trapiche-Empozada formations, Fig. 3) still show very restricted areal distribution and limestone conglomerate mega-beds (Astini 1994). These indicate important block faulting with local exposure of limestone cores that provided the intra-basin source for the megabreccias. Different uppermost Ordovician (postHirnantian) and Silurian successions (Fig. 3) overlap, in most of the central Precordillera, different Ordovician substrates (Astini & Maretto 1996). This would constitute the first clear overlap of Late Ordovician extension although, as suggested by Astini & Maretto (1996), seismicity related to extension faulting would have continued during most of the Silurian.

Discussion The selection of an evolutionary model of the Precordillera must rely entirely on evaluation of the available stratigraphical and associated geophysical and geochemical data. Although the microcontinental model does not allow precise relative longitudinal positioning of Laurentia and Gondwana, it is well supported by the stratigraphical data discussed in this and previous papers. Furthermore, indirect positioning can be obtained on reasonable bases as discussed by Thomas & Astini (1996a). As stressed throughout this paper, the geological evidence favours Early Cambrian rifting and progressive isolation of the Precordilleran microcontinent

25

during the drifting stage. This is more consistent with the traditional plate palaeogeography, where the Iapetus Ocean was the epicentre of intense 'traffic' of terranes and island-arcs (one of which was the Precordillera) in a complex mosaic that resembles the present-day SW Pacific (e.g., MacNiocaill et al. 1997). It is important to note that considering the Precordillera terrane as an independent microcontinent does not preclude the 'end-run' hypothesis proposed by Dalziel (1991) which seems to correspond to the slightly younger history of Laurentia (see Meert et al. 1994; Torsvik et al. 1996). Denying the tectonic tracer hypothesis for the Precordillera unfortunately implies that longitudinal positioning of Laurentia during its relative drifting toward the north remains unconstrained. According to present stratigraphical knowledge, while Laurentia remained within the tropics, the Precordillera terrane drifted southward and by early Late Ordovician (a time when the Iapetus Ocean was fully developed) it occupied a relatively high southern latitude (Astini 1995a). Convergence between Laurentia and Gondwana to form the Taconic and Ocloyic (Famatinian) orogenies (Dalla Salda et al. 1992b), respectively, in South and North America, is presently unsupported. The Precordillera was most probably isolated from Laurentia by a western portion of the Iapetus Ocean during the Taconic Orogeny, the latter having its origin in different terranes that accreted against Laurentia as part of the complex mosaic of terranes and island arcs that characterized the Iapetus Ocean (Fig. 10). The Ocloyic orogeny can, in turn, be considered as the product of collision of island-arc complexes, such as the Famatina System and the Puna (Conti et al. 1996), and the Precordillera terrane against the western margin of Gondwana (Fig. 10). Early Cambrian rifting of the Precordillera block, at variance with the 'bridge' model, explains the development of graben and grabenfilled successions and the evolution of the passive margin stage. Mafic rocks from an ophiolite section proved older than Caradoc have recently been found in the western Precordillera (Davis et al. in press). Pods of serpentinite and metagabbro within the Ouachita thrust belt/ accretionary prism in the Laurentian margin suggest oceanic crust beneath the Late Cambrian and younger slope-and-rise deposits of the Ouachita embayment (Nielsen et al. 1989; Viele & Thomas 1989). Additionally, faunal assemblages diagnostic of shelf margin-slope biofacies (Palmer et al. 1996) were found in the Middle Cambrian olistoliths of the western Precordillera

26

R . A . ASTINI

(Bathyurids)

. - . ........ ,(~

.

....... "--.

..

~

,

_

30.~

shales and fine-grained siliciclastic rocks . . . . . . . .

.

,

carbonates low-latitude

.

~

volcaniclastic and

volcanic rocks

.-'"

(B

'~

shallow-marine siliciclastic rocks BA

/

_~-. . . . . . . . . . . . . . ....0-Oo ~

"".4Y

/'

\

. ~//(Northern)

/

Lt:~#' "~ g .~, / ~"

9 . . . ~ ~ut~s*

\

"" i,

w-'"" lll'"..~,

co

~

| SOUTH POLE

d

i

i

i

! /'

~,J-

..

.J

/

,/

-'//

/' //'

....,

/

,/ /

~RECORDI TERRANE LLERA,~, ~"'"", ! (mixed Toquima-Table

(Bath~iurids) + enderiti~s

/

/

/

(Southern),i

Head and'Celtic genera)

/

-'0 ~

/

y/

FAMATINA ' ISLAND ARC

(mixed Celtic and ".Mediterranean genera)

c. 470 Ma

,/

/ ,+'/

~N l

.."

Fig. 10. Early Ordovician palaeogeographic reconstruction showing distribution of brachiopod faunas, trilobites and facies belts. L, Laurentia; S, Siberia; B, Baltica; G, Gondwana; NCH, North China block; SCH, South China block; AM, Armorica; BM, Bohemia; AN, Antarctica; AUST, Australia; IN, India. Northern and Southern Iapetus are indicated. (modified from Astini et al. 1995a; Benedetto & S{mchez 1996; Torsvik et al. 1996; Mac Niocaill et al. 1997).

(Vaccari & Bordonaro 1993). In the absence of an ophiolite suite in structurally complicated areas, Fortey & Cocks (1986) have suggested that marginal faunal belts or platform edge biofacies can be used, in principle, to indicate the proximity of former ocean basins. These findings reinforce the suggested Early Cambrian rifting phase and expansion of an oceanic basin between the Precordillera and the Texas promontory, which would further develop into a true barrier between the Precordillera and Laurentia. A possible modern analogue of the Precordillera-microcontinental model, where the jump of

a spreading ridge occurs inboard into continental crust to separate a smaller crustal block, is the Jan Mayen Ridge. This ridge is a microcontinent rifted from east Greenland, with transform motion on the Jan Mayen fracture zone (Eldholm e t al. 1990; Thomas in press). Dalziel (1997) pointed out that offset lengths on transform segments usually shorten rather than lengthen in present-day examples and, hence, this would constitute a serious objection to the microcontinent model. Nevertheless, Vink e t al. (1984) have described several cases of rifting of small continental fragments or microplates

THE LAURENTIAN PRECORDILLERA TERRANE where inboard ridge-jumps and lengthening offsets on transform segments have preferentially occurred. Major elements that collectively support the rifting, drifting, and collisional model are: (1) rifted margin stratigraphy, (2) development of a typical thermo-tectonically controlled passive-margin carbonate bank, (3) progressive differentiation in faunas and lithology, (4) apparent isolation of the carbonate platform through most of the Tremadoc and Arenig (basically much of the North American Ibexian), (5) beginning of widespread drowning together with inflow of new faunas and general change toward cooler water environments, (6) ash beds reflecting progressive the proximity of volcanic arcs, (7) pervasive instability throughout the basin (suggested from stratigraphical architecture and irregular distribution of sediments and hiatuses), and (8) direct glacial evidence related to Hirnantia fauna. This evidence means that the Precordillera could be considered one of the best constrained allochthonous terranes in Earth history.

Conclusions (1) Based on a reappraisal of the available stratigraphical evidence it is suggested that the Precordillera terrane should not be considered a tectonic tracer, but should instead be categorized as an independent lithospheric fragment exotic to Gondwana. (2) Passive margin stratigraphy, subsidence analysis and marginal faunal belts, associated with a straightforward west-polarity, indicate a fully developed rift stage during the Early Cambrian. (3) A mixed succession of evaporites, red siliciclastic rocks and dolomites underlying the thick Cambrian to Lower Ordovician carbonate platform record a syn-rift, restricted circulation, low latitude, divergent margin setting. Rift to drift completion can be deduced from the overlap of the syn-rift facies by transgressive glauconitic sandstones and limestones with the first open-marine faunas, of late Early Cambrian age. (4) The Precordillera block has its source in the Ouachita embayment in southeastern Laurentia. Stratigraphical and kinematic constraints favour an asymmetric rifting model for the separation of Precordillera and the Texas promontory. (5) After the Late Cambrian, the Precordillera block apparently moved beyond the Alabama promontory, breaking all connections with

27

Laurentia. Progressive separation of the Precordillera block from Laurentian during the drifting stage is mainly supported by the changeover in faunas, although progressive isolation is also indicated by sedimentary features and exponential decay of thermal history. (6) Gradual proximity to Gondwana and bordering island- or volcanic-arcs is marked by prominent ash layers included in the upper part of the Lower Ordovician limestones and a forebulge flexural stage inferred by progressive drowning of the carbonate platform. (7) Mid- to Late Ordovician widespread extension is deduced from the complex architecture and the development of local hiatuses, and restricted depocentres with dominantly siliciclastic sedimentation and scarce carbonate remnants in the Precordillera. This second extension episode is separated by more than 70 Ma from the initial rifting of the Precordillera from Laurentia, and is related to post-collisional disruption and major change in plate motions following collision. (8) Late Ordovician glacial deposits typical of Gondwana overlap the Precordillera region, yielding a minimum age for the proposed accretion of the Precordillera terrane, although pervasive indirect sedimentary evidence of extensional tectonics by Mid-Ordovician allows bracketing the collision near the early MidOrdovician. (9) Stratigraphical evidence points toward different timing and nature of the Ocloyic and Taconic orogenies in South and North America, respectively, although both are related to the collision of terranes and island arcs on opposite margins of the evolving Iapetus. Various Argentine institutions supported my research in the Argentine Precordillera: Consejo Nacional de Investigaciones Cientificas y T+cnicas (CONICET), Consejo de Investigaciones Cientificas y T~cnicas de la Provincia de C6rdoba (CONICOR), Fundaci6n Antorchas, and Secretaria de Ciencia y T~cnica de la Universidad Nacional de C6rdoba (SECyT). Support from the University of Western Michigan and the University of Texas at Austin allowed me to visit type Laurentian localities. I am grateful to Bob Hatcher Jr., for patient advice on a draft manuscript and suggestions that greatly improved this work, as did constructive reviews by John Cooper and David Macdonald. I very much appreciate continuous discussion of Precordilleran topics with colleagues in C6rdoba, especially Luis Benedetto, Marcelo Carrera, Edsel Brussa, Emilio Vaccari and Fernando Cafias, and I also wish to thank Augusto Rapalini, Victor Ramos, Bill Thomas, Steven Davis, Sarah Roeske, Warren Huff, Stig Bergstr6m, Chuck Mitchell, Carlos Cingolani, Christopher Schmidt and Ian Dalziel.

28

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e

f

R.A.

e

r

e

n

c

e

s

ALBANESI, G. L. & BARNES, C. R. 1997. The Middle Ordovician conodont Paroistodus horridus in the Argentine Precordillera. 6th International Conodont Symposium, Warsaw, Abstracts, 1. , HONICKEN, M. & ORTEGA, G. 1990. Amorphognathus supervus (Conodonta) from Trapiche Formation (Upper Ordovician), Cerro Potrerillos, J/tchal Department, San Juan Province, Argentina. First Latin American Conodont Symposium, Academia de Ciencias, C6rdoba, Abstracts, 109-110. AITKEN, J. D. 1966. Middle Cambrian to Middle Ordovician cyclic sedimentation, southern Rocky Mountains of Alberta. Bulletin of Canadian Petroleum Geology, 14, 405-441. ALLMENDINGER, R. W., FIGUEROA, D., SYDNER, D., BEER, J., MPODOZIS, C. & ISACKS, B. L. 1990. Foreland shortening and crustal balancing in the Andes at 30~ latitude. Tectonics, 9, 789-809. ARMELLA, C. 1994. Thrombolitic-stromatolitic cycles of the Cambro-Ordovician boundary sequence, Precordillera oriental basin, western Argentina. In: BERTRAND SARFATI, J. & MONTY, C. (eds) Phanerozoic Stromatolites, 2, Kluwer Academic Publishers, Netherlands, 421-441. ASTINI, R. A. 1991. Paleoambientes sedimentarios y secuencias depositacionales del Ordovicico cldstico de la Precordillera Argentina. PhD thesis, Universidad Nacional de C6rdoba. - - 1 9 9 2 . Tectofacies ordovicicas y evoluci6n de la cuenca eopaleozoica de la Precordillera Argentina. Estudios Geol6gicos, 48, 315-327. - - 1 9 9 4 . Las megaturbiditas de la Formaci6n Trapiche (Ordovicico superior de la PrecordiUera): procesos sedimentarios y marco geol6gico. Quinta Reunidn Argentina de Sedimentologia, San Miguel de Tucumfin, Argentina, 1, 107-112. - - 1 9 9 5 a . Paleoclimates and paleogeographic paths of the Argentine Precordillera during the Ordovician: evidence from climatically sensitive lithofacies. In: COOPER, J. D., DROSER, M. L. & FINNEY, S. C. (eds) Ordovician Odyssey. Society of Economic Paleontologists and Mineralogists, Book 77, 177-180. --1995b. Geologic meaning of Arenig-Llanvirn diachronous black shales (Gualcamayo Alloformation) in the Argentine Precordillera, tectonic or eustatic? In: COOPER, J. D., DROSER, M. L. & FINNEY, S. C. (eds) Ordovician Odyssey. Society of Economic Paleontologists and Mineralogists, Book 77, 217-220. - - 1 9 9 5 c . Sedimentologia de la Formaci6n Las Aguaditas (talud carbon/~tico) e implicancias estratigr/lficas en la cuenca precordillerana oriental durante el Ordovicico. Revista de la Asociacidn Geoldgica Argentina, 50, 143-164. --1996. Las fases diastr6ficas del Paleozoico medio en la Precordillera del oeste argentinoevidencias estratigraficas-. XIII Congreso Geoldgico Argentino y III Congreso de Exploracidn de Hidrocarburos, Buenos Aires, Actas, V, 509-526.

ASTINI - - 1 9 9 7 a . Las unidades calc~treas del Ordovicico Medio y Superior de la Precordillera Argentina como indicadores de una etapa extensional. Segundas Jornadas de Geologia de Precordillera, San Juan, Argentina, 1, 8-14. 1997b. Stratigraphic evidence of two-stage rifting and collison in the Laurentian derived Precordillera terrane, south-central Andes. Geological Society of America, Annual Meeting, Abstracts with Programs, 29, A-116. in press. El "Conglomerado de Las Vacas" y el Grupo Trapiche de la Precordillera: Tectbnica distensiva en el Ordovicico Superior. Revista de la Asociacidn Geoldgica Argentina. & BENEDETTO, J. L. 1996. Tectonostratigraphic development and history of an allochthonous terrane in the Pre-Andean Gondwana margin: the Argentine Precordillera. 3rd International Symposium on Andean Geodynamics, St Malo, France, Extended Abstracts, 759-762. & BRUSSA, E. D. 1997. Dos nuevas localidades fosiliferas en el Conglomerado de Las Vacas (Caradociano) en la Precordillera Argentina: Importancia cronoestratigrfifica. Reuni6n de Comunicaciones Paleontol6gicas. Ameghiniana, 43, 231. & CA~AS, F. L. 1995. La Formaci6n Sassito, una nueva unidad calc/~rea en la Precordillera de San Juan: sedimentologia y significado estratigrfifico. Asociacidn Argentina de Sedimentologla, Revista, 2, 19-37. & MARETTO, H. M. 1996. Anfilisis estratigr/lfico del Silflrico de la Precordillera Central de San Juan y consideraciones sobre la evoluci6n de la cuenca. XIII Congreso Geolbgico Argentino y III Congreso de Exploracibn de Hidrocarburos, Buenos Aires, Actas, I, 351-368. • VACCARI,N. E. 1996. Sucesi6n evaporitica del Cfimbrico inferior de la Precordillera: significado geol6gico. Revista de la Asoeiacidn Geoldgica Argentina, 51, 9%106. - - , BENEDETTO,J. L., & VACCARI, N. E. 1995a. The Early Paleozoic evolution of the Argentine Precordillera as a Laurentian rifted, drifted and collided terrane: a geodynamic model. Geological Society of America Bulletin, 107, 253-273. , RAMOS, V. A., BENEDETTO, J. L., VACCARI, N. E. & CAlqAS, F. L. 1996. La Precordillera: un terreno exotico a Gondwana. XIII Congreso Geoldgico Argentino y III Congreso de Exploracidn de Hidrocarburos, Buenos Aires, Actas, V, 293-324. - - - , VACCARI, N. E., THOMAS, W. A., RAYMOND, D. E. & OSBORNE, W. E. 1995b. Shared evolution of the Argentine Precordillerra and the Southern Appalachians during the Lower Cambrian. Geological Society of America, Abstracts with Programs, 27, 458. BALDIS, B. A. & BLASCO, G. 1973. Trilobites Ordovicicos de Pon6n Trehue, Sierra Pintada de San Rafael, provincia de Mendoza. Ameghiniana, 10, 72-88. & BORDONARO, O. 1981. Evoluci6n de las facies carbon/lticas en la cuenca C~tmbrica de la Precordillerra de San Juan. VIII Congreso Geoldgico Argentino, San Luis, Actas, II, 385-397. -

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THE LAURENTIAN PRECORDILLERA TERRANE --

& 1982. Comparaci6n entre el Cfimbrico de la "Great Basin" norteamericana y la Precordillera de San Juan, Argentina, su implicancia intercontinental. V Congreso Latinoamericano de Geologia, Buenos Aires, 1, 97-108. & 1985. Variaciones de facies en la cuenca cfimbrica de la Precordillera Argentina, y su relaci6n con la generacidn del borde continental. VI Congreso Latinoamerieano de Geologia, Bogotfi, Actas, 1, 149-161. , ARMELLA, C. & CABALERI, N. 1985. Desarrollo de la plataforma carbonfitica Ordovicica argentina. VI Congreso Latinoamericano de Geologia, Bogota, Actas, 1, 165-174. BARNES, V. E., BELL, W. C., CLABAUGH,S. E., CLOUD, P. E., JR., MCGEHEE, R. V., RODDA, P. U. & YOUNG, K. 1972. Geology of the Llano region and Austin area, field excursion. Guidebook No. 13. Texas Bureau of Economic Geology, University of Texas, Austin. BENEDETTO, J. L. 1993. La hip6tesis de la aloctonia de la Precordillera Argentina: un test estratigr/tfico y biogeogrfifico. XII Congreso Geoldgico Argentino y H Congreso de Exploracion de Hidrocarburos, Mendoza, Actas, III, 375-384. - - 1 9 9 8 . Early Palaeozoic brachiopods and associated shelly faunas from western Gondwana: their bearing on the geodynamic history of the pre-Andean margin. This volume. -& ASTINI, R. A. 1993. A collisional model for the Early Paleozoic stratigraphic evolution of the Argentine Precordillera. 2nd International Symposium on Andean Geodynamics, Oxford, Extended Abstracts, 501-504. -& HERRERA, Z. A. 1987. Primer hallazgo de braqui6podos y trilobites de la Formaci6n Trapiche (Ordovicico Tardio), Precordillera Argentina. X Congreso Geol6gico Argentino, Tucumfin, Actas, III, 73-76. & SANCHEZ, T. M. 1996. Paleobiogeography of brachiopod and molluscan faunas along the South American margin of Gondwana during the Ordovician. In: BALDIS, B. A. & ACElqOLAZA, F. G. (eds) Early Paleozoic Evolution in NW Gondwana. Serie Correlaci6n Geol6gica, Universidad Nacional de Tucum/m, 12, 23-38. & VACCARI,N. E. 1992. Significado estratigrfifico y tect6nico de los complejos de bloques cambroordovicicos resedimentados de la Precordillera Occidental, Argentina. Estudios Geol6gicos, 48, 305-313. - - , CARRERA, M. G. & S,h,NCHEZ, T. M. 1995. ~Fhe evolution of faunal provincialism in the Argentine Precordillera during the Ordovician: new evidence and paleontological implications. In: COOPER, J. D., DROSER, M. L. & FINNEY, S. C. (eds) Ordovician Odyssey. Society of Economic Paleontologists and Mineralogists, Pacific Section, Book 77, 181-184. BERNOULLI, D. & JENKYNS, H. 1974. Alpine, Mediterranean, and central Atlantic Mesozoic facies in relation to the early evolution of the Tethys. In: DOTT, R. H. & SHAVER, R. H. (eds) Modern and ancient geosyndinal evolution. Society of -

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The stratigraphical reeord of the Argentine Preeordillera and its plate-tectonic background MARTIN

KELLER,

WERNER

BUGGISCH

& OLIVER

LEHNERT

Institut ffir Geologie und Mineralogie, Universitdit Erlangen, Schlossgarten 5, D-91054 Erlangen, G e r m a n y Abstract: The stratigraphical record of the Argentine Precordillera from Early Cambrian to

Late Devonian times reveals its plate-tectonic history from incipient rifting and the evolution of a marginal platform to the separation from Laurentia, its drift in higher latitudes and amalgamation with Gondwana. This pre-Carboniferous succession can be subdivided into four supersequences bounded by major unconformities and their main features are discussed with respect to plate-tectonic implications. The basal two supersequences which include the carbonate platform deposits are subdivided into 13 third-order sequences, each with a duration of 2-10 Ma. Supersequence A reflects intracratonic rifting, creating a graben system and forming a marginal plateau. In its in higher part a progradational carbonate complex covered the entire platform. Supersequence B shows the development of an aggradational carbonate succession and the evolution of reef ecosystems comparable to those that developed around the margins of the Ouachita embayment. It also shows the demise of the carbonate platform by drowning. Deposits of Supersequence C reflect crustal extension and rifting, which led to the final separation of the Precordillera from mainland Laurentia. Supersequence D reveals the approach, and probably the accretion, of the Argentine Precordillera to Gondwana.

The Argentine Precordillera (AP) is a morphostructural entity between the main Andean Cordillera to the west and the Sierras Pampeanas to the east (Fig. 1). It differs from the surrounding geological provinces in exposing a thick Cambro-Ordovician carbonate platform succession. Within this succession, there is a Cambrian trilobite fauna typical of the Early Palaeozoic margins of Laurentia (Borrello 1971; Vaccari 1995). Recent investigations have shown that not only the trilobite fauna is similar to that of Laurentia, but also the Late Cambrian and Early Ordovician conodont faunas (Lehnert 1995). In addition, there is a close similarity between reef ecosystems in the AP and those around the Ouachita embayment along the southern margin of Laurentia (Keller & Flfigel 1996; Keller 1997). The uniqueness of the carbonate succession, together with the palaeontological data and isotope data from the presumed basement of the AP, led to speculation that the AP was a terrane exotic to South America, derived from Laurentia (Bond et al. 1984; Keppie 1991; Dalla Salda et al. 1992a, b, 1993; Astini et al. 1995; 1996, Dalziel et al. 1994; Dalziel 1997). Although there are models which try to interpret the origin of the AP as being autochthonous (Gonzfilez Bonorino & Gonzf, lez Bonorino 1991) or parautochthonous (Baldis et al. 1989; Loske 1992), there has been increasing agreement that it is

indeed a Laurentia-derived terrane (see Dalziel 1997 and Dalziel et al. 1996 for discussion). However, models explaining the provenance, the timing of the possible transfer, and the accretion of the AP to Gondwana are highly controversial (Dalla Salda et al. 1992a; Dalziel et al. 1994; Astini et al. 1995, 1996; Keller & Dickerson 1996; Keller 1997). Although many authors believe in Mid-Ordovician accretion, we take the Carboniferous deposits of the Paganzo basin, which unconformably rest on the AP and adjacent terranes, as the first unequivocal sign that the AP had finally been accreted to Gondwana. Based on this minimum age of accretion, we will discuss the pre-Carboniferous sediments of the AP and their depositional environment in order to interpret the geotectonic background in which they formed. Many of our conclusions are drawn under the assumption that the AP is indeed a Laurentia-derived terrane. We will discuss the most important lithostratigraphical units and make some reference to less important sediments, but will not attempt to give a complete stratigraphy of the AP. Similarly, we will only briefly mention the most important sequence-stratigraphical features of the thirdorder cycles in the carbonate platform deposits, i.e. those that help in elucidating the platetectonic history. For a more detailed sequencestratigraphical discussion of the carbonates the reader is referred to Keller (1997).

KELLER, M., BUGGISCH,W. & LEHNERT,O. 1998. The stratigraphical record of the Argentine Precordillera and its plate-tectonic background. In: PANK~tJRST, R. J. & RAPELA, C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 35-56.

36

M. KELLER E T A L .

Fig. 1. Morpho-structural units of NW Argentina and adjacent areas and components of the Cuyania terrane (1, Eastern Sierras Pampeanas; 2, Sierra de Famatina; 3, Western Sierras Pampeanas; 4, valley of Iglesias-Calingasta-Uspallata).

Regional overview Geologically, the AP constitutes a high-level fold and thrust belt of mainly Miocene age. The sedimentary succession starts with Lower Cambrian sediments and, with several gaps, continues into the Triassic. Younger Mesozoic deposits are absent but, with the onset of crustal shortening and the corresponding uplift of the Andes, a thick succession of Tertiary and, locally, Quaternary sediments was deposited in terrestrial environments. The Cambrian to lower Middle Ordovician sediments are predominantly carbonates, whereas in younger deposits terrigenous clastic rocks prevail. The contact with the underlying basement is nowhere exposed, but there is indirect evidence for a metamorphic basement from xenoliths within Tertiary volcanic rocks (Abbruzzi et al. 1993) and from basement clasts in the Ordovician continental margin facies (Keller 1995). Isotope data from the xenoliths show values typical of the Grenville-age belts along the Appalachian margin of Laurentia. Similar

Fig. 2. Distribution of Cambro-Ordovician sediments in the Argentine Precordillera (modified from Keller & Bordonaro 1993) and localities mentioned in the text. values have been obtained recently from basement rocks in the western Sierras Pampeanas (McDonough et al. 1993; Kay et al. 1996). The recognition of Grenville-type rocks both beneath the AP and in the western Sierras Pampeanas (Sierra Pie de Palo, Fig. 1) led Ramos (1995) to propose the existence of a 'Cuyania' terrane composed of Grenvillian crust and a carbonate platform cover with Laurentian faunas. An additional fragment of the Cuyania terrane has been identified near San Rafael (Bordonaro et al. 1996; Fig. 1), where Lower Ordovician carbonates with a Laurentian fauna are exposed.

PRECORDILLERA STRATIGRAPHY Within the AP proper (Fig. 2), two fundamentally different tectono-sedimentary environments are recognized (e.g., Baldis et al. 1982; Astini 1992; Keller 1997): a carbonate platform of Cambrian to early Middle Ordovician age overlain by predominantly siliciclastic platform deposits (hereafter referred to as the platform or former platform area), and a western siliciclastic basin which developed mainly during the

37

Mid- and Late Ordovician as a response to rifting (here referred to as the western basin).

Stratigraphy The pre-Carboniferous sedimentary succession of the AP (Fig. 3) can be subdivided into several major, unconformity-bounded supersequences:

Fig. 3. Lithostratigraphy of pre-Carboniferous rocks in the AP. This chart is a compilation of published information from many different sources.

M. KELLER ET AL.

38

Supersequence A (Late Early Cambrian-Late Cambrian), Supersequence B (Early Ordovicianearliest Mid-Ordovician), Supersequence C (MidOrdovician-Late Ordovician), and Supersequence D (Silurian-Late Devonian). Supersequences A and B are composed of a total of 13 sequences (Fig. 4) spanning the Mid-Cambrian to earliest Mid-Ordovician. The duration of each sequence varies between 2 M a and 10Ma. This is within the average duration of third-order cycles (1-10Ma) as originally described by Vail et al. (1977). Some of these sequences are composed of several smaller-scale sequences, which from their size and average duration are intermediate between

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The Cerro Totora Formation consists of marginal-marine fine-grained sandstones and siltstones in the lower part of the succession. Up section, these grade into evaporites with carbonate interbeds, which in turn are overlain by sandstones, siltstones, and shales alternating with oolitic and bioclastic grainstones. Locally, quartz arenites are present near the top of the formation. The depositional environment has been interpreted as peritidal marine (Astini & Vaccari 1996). Frequent periods of restricted conditions led to the formation of evaporites in the middle part of the succession. In contrast, the presence of Early Cambrian olenellid trilobites (Vaccari 1994) in the upper part indicates normal marine salinity.

La Laja Formation (>525m)

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Strata of Supersequence A are exposed in the Guandacol area, in the Sierra Chica de Zonda, and in the Sierra de Villicum. Four formations have been recognized in the carbonate platform deposits of the AP. In addition, deep-water limestones of Middle and Late Cambrian age (Bordonaro & Banchig 1996) are preserved as olistoliths in Ordovician continental margin sediments.

Lithostratigraphy.

1500m

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Cerro Totora Formation (c. 340m)

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the third-order sequences and the typical fifthorder small-scale cycles with thicknesses varying between 1 m and 5 m. Consequently, these intermediate-scale cycles are regarded as fourthorder cycles.

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] ] 0m Fig. 4. Third-order sequences of the Argentine Precordillera carbonate platform succession and their thicknesses. Note increasing cycle thickness towards the Cambrian-Ordovician boundary (sequence boundary 9) and decreasing thicknesses thereafter.

Lithostratigraphy. The La Laja Formation is the oldest carbonate platform unit preserved in the AP. It displays a varied spectrum of carbonate rocks with intercalated intervals of siliciclastic deposits. The sediments have been grouped into depositional assemblages (Keller & Bordonaro in press) or facies associations (Keller 1997). The lime mudstone-wackestone-intraclast packstone association is dominated by muddy textures containing well-preserved fossils, among which are trilobites, hyolithids, rare phosphatic brachiopods, and sponge spicules. Some of the mudstones are strongly burrowed. Intraclast packstones are graded and in places show an erosional contact with the underlying

PRECORDILLERA STRATIGRAPHY beds. The sediments were deposited below fairweather wave base in a well-aerated environment. The intraclast packstones represent tempestites, their clast content was derived from shallow shelf areas. The oncolite-wackestone association is composed of thick, in places cross-bedded, oncolite beds intercalated in wackestones with trilobites, hyolithids, and brachiopods. Erosional features at the base of the oncolites are present locally. The depositional environment of this association is similar to that of the lime mudstone-wackestoneintraclast packstone association. The oncolites are also interpreted as tempestites, although their source area was different. The original site of oncoid formation has not been found. The oolite-oncolite association is composed of oolitic and/or oncolitic packstones and grainstones. Many of the horizons show crossbedding, and locally well-developed herringbone cross-stratification has been observed. Similar sedimentary features were found in the packstone-grainstone association, but these rocks are predominantly composed of fossil hash (trilobites, hyolithids, eocrinoids, brachiopods). Detrital quartz is present is some beds. Sediments of both associations represent winnowed sand bars, barriers, and subtidal shelf sands. The depositional environment was shallow subtidal and lower intertidal where wave and current activities constantly removed the lime mud and led to the formation of herringbone cross-stratification. The siltstone association is easily recognized in the field by its yellow and brown colours which contrast with the dark grey and black colours of the limestones. Besides siltstones, silty marlstones, marlstones, eocrinoidal or oolitic grainstones are part of this association. Wave ripples and linsen and flaser bedding are the most prominent bedforms. The sandstonegrainstone association is composed of bioclastic, in places oolitic, grainstones and thin beds of quartz arenites or calcareous sandstones. Coquinas of olenellids are common. In both associations, lithology in combination with sedimentary structures indicates a shallow subtidal shelf with a strong tidal influence. However, in three sections thick quartz arenites alternating with black shales are present. Tectonic deformation has obliterated all sedimentary structures; consequently, an interpretation of this succession is difficult. Sediments of the La Laja Formation represent three depositional systems: a carbonate shoal complex with dominantly grainstones and packstones, an open marine environment with mudstones and wackestones, and a shallow shelf association influenced by terrigenous detritus.

39

Mudstones and wackestones are interpreted to represent both back-barrier, lagoon deposits and sediments which accumulated seaward of the main high energy zone.

Sequence stratigraphy. Two types of shallowing-upward successions have been observed. There are calcareous successions with a thickness between some tens of metres and more than 100m. The successions start with sediments of the lime mudstone-wackestone-intraclast packstone association. As this association represents the deepest environment present in the La Laja Formation, the basal bounding surface is always a flooding surface. Up section, bioclastic packstones and grainstones are developed. Oolites or oncolites form the top of the carbonate succession. Within the individual successions no smallscale cycles have been observed. Several of these cycles are stacked to form larger-scale sequences. Four of the latter within the La Laja Formation (Fig. 4) span approximately 10 Ma, hence they have to be regarded as third-order cycles. The most characteristic feature of these cycles are basal units composed of the siltstone association. No dramatic environmental change is indicated by the switch from carbonate production at the top of a carbonate cycle to siliciclastic deposition of the siltstone associations. In contrast, the upper boundary of the terrigenous units is always sharp and records important flooding. Consequently, the surface separating the siltstone units from the overlying carbonate intervals is interpreted as the transgressive surface (Keller 1997) which marks the beginning of the transgressive systems tract. The transgressive systems tract is relatively thin and is composed mainly of rocks with muddy textures. There is no distinct surface which might be interpreted as a maximum flooding surface. Nevertheless, the change from mud-dominated textures in the lower part of the succession to grain-supported textures in the upper part seems to reflect the transition from the transgressive systems tract to the highstand systems tract. The relatively thick highstand systems tract together with the internal architecture of the entire succession mark a catch-up system (Kendall & Schlager 1981). In such systems, shoal deposition (oolite-oncolite association in the La Laja Formation) is restricted to the late highstand. In most sequences within the La Laja Formation, sediments of the oolite-oncolite association were deposited during this interval related to sealevel stillstand or relative sea-level fall. The position of the siltstone units between late highstand deposits and the transgressive surface at the base of the subsequent carbonate interval

40

M. KELLER ET AL.

indicates that the siltstone units are related to a sea-level lowstand and that their lower boundary is the sequence boundary. No signs of subaerial exposure have been observed along these contacts, nor is there a basinward shift in facies. Consequently, the sequence boundaries are type-2 boundaries, and the siltstone intervals represent the corresponding shelfmargin systems tracts.

The Zonda and La Flecha Formations Lithostratigraphy. Both formations are composed predominantly of peritidal dolomites. Secondary dolomitization has obliterated most of the primary structures in the Zonda Formation (200-300 m). Where preserved, sedimentary structures are comparable to those of the La Flecha Formation. The sediments of the latter (400-700m) have been described in detail by Keller et al. (1989), Cafias (1995a), and Armella (1989a, b). Rocks of the mudstone-wackestone association contain a sparse fauna (trilobites, hyolithids) or are bioturbated. The sediments were deposited in shallow subtidal areas, probably under restricted conditions. A few intercalated trilobite packstones represent episodic storms. In the intraclast grainstone-oolite-thrombolite association, the intraclasts reflect erosional processes that affected the entire spectrum of rocks present in both formations. Oolite grainstones are present as continuous beds, often with well developed herringbone cross-stratification. However, they also fill erosional channels which in places cut down almost 2m into thrombolite mounds. Locally, the oolites onlap individual mounds demonstrating a close relation between mound growth and oolite deposition. The individual thrombolites either grew as isolated mounds or they form laterally linked structures several metres high and hundreds of metres wide. This association records complex sediment build-up interactions in shallow subtidal and lower intertidal areas. The stromatolite association is composed of different types of stromatolites and accompanying mudstones and grainstones. LLH-stromatolites are most abundant, but SH morphotypes have also been observed. The association formed in environments extending from the shallow subtidal to the lowermost supratidal. In the microbial laminite-breccia association, thick mudstone intervals, in places with evaporite pseudomorphs, represent storm events during which lime mud was transported from its original site of formation, mainly the shallow

subtidal, onto the supratidal flats. Mudcracks, tepee structures, and pseudomorphs after evaporites testify to prolonged episodes of subaerial exposure, desiccation, and evaporation. The association described here was mainly formed in higher intertidal and supratidal environments. Finally, there are some mudstones with large spheroidal mudlayers and abundant microscopic cracks. These have been interpreted as calcrete horizons (Keller et al. 1989) and were attributed to a terrestrial setting. Rocks of both formations were deposited in a peritidal environment, many of them under restricted, hypersaline conditions. Small-scale (fifth-order) cycles are a prominent feature in all sections, their internal architecture, however, does not permit a distinction of more proximal from more distal cycles. The only evidence of somewhat 'deeper' environments is present in the Guandacol area, where there is a higher percentage of shallow-subtidal mudstones in the cycles (Cafias 1995a).

Sequence stratigraphy. Sedimentologically, the lower part of the Zonda Formation is a continuation of the uppermost sequence of the La Laja Formation (sequence 5; Fig. 4). This is indicated by the transition from oolites with herringbone cross-stratification (uppermost part of the La Laja Formation) to inter- and supratidal dolostones of the basal Zonda Formation. These rocks are abruptly overlain by dark subtidal mudstones at the base of another major shallowing-upward succession. The upper boundary of this cycle coincides with the boundary between the Zonda and La Flecha Formations. The La Flecha Formation is composed of two similar shallowing-upward sequences (sequences 7 and 8; Fig. 4), but there is no well-defined boundary between them. Within the lower cycle the abundance of calcrete horizons increases towards the top but calcretes are absent above the presumed cycle boundary. A reversed pattern is visible in the distribution of thrombolites: they become less abundant towards the top of the cycle, are absent in the uppermost interval, but regain importance at the base of the next cycle. All three sequence boundaries within the Upper Cambrian strata either show signs of subaerial erosion, coarse detrital quartz, abundant evaporites, or concentrations of calcrete horizons just beneath the main surface. Consequently, each of these sequence boundaries has to be regarded as a type-1 sequence boundary (Keller 1997). These sequences were deposited during approximately 10Ma, which qualifies them as third-order sequences.

PRECORDILLERA STRATIGRAPHY

Middle and Upper Cambrian deep-water limestones and marlstones These rocks, which recently have been combined under the informal name of the 'La Cruz limestones' (Keller 1997), are mainly present as olistoliths within the Ordovician continental margin facies. Only the outcrop at Cerro Pelado shows these deep-water limestones in a presumably autochthonous position above peritidal dolomites of the La Flecha Formation. The limestones and marlstones are dark grey and fine grained, and thin- to medium-bedded. Shales have been observed only locally. The fauna consists of agnostid trilobites and sponge spicules. Graded beds, in places with erosional structures at the base, and slumped beds are also present. In general, the components are less than 1 mm in diameter and composed of platformderived material. The horizons described here are interpreted as distal tempestites, deposited in a background environment of deep-water limestones and marlstones. It is noteworthy that wherever these rocks are exposed as olistoliths within the continental slope facies, there are no indications of mass-flow deposits which might indicate the existence of a Mid- or Late Cambrian platform margin or slope. In contrast, we believe that the platform gradually passed into deep-water environments without a pronounced slope. Platform configuration might have been a homoclinal ramp (Read 1982, 1985; Burchette & Wright 1992). Within the most important olistolith, at the Los Tuneles section (Fig. 2), an unconformity is present separating Middle Cambrian deep-water limestones from latest Cambrian or even lowermost Ordovician slope deposits (Keller 1995). These are the oldest slope sediments hitherto known from the AP.

Evolution of Supersequence A The basal succession of supersequence A consists of redbeds alternating with evaporite layers (Cerro Totora Formation). Astini et al. (1995) interpreted this succession as a rift-related sequence indicating the separation of the AP from Laurentia. In our view, its localized presence and relatively reduced thickness, together with the absence of any indications of riftrelated volcanism or ocean floor, all point to intracontinental rifting similar to that of the coeval Birmingham graben of Laurentia (Thomas 1991). The uppermost horizons of the Cerro Totora Formation are coeval to the basal beds of

41

the La Laja Formation (El Estero Member). In this part of the succession, both formations show an alternation of sandstones, shales, and grainstones which constitutes the transition to carbonate platform sedimentation. Cafias (1988) described a hardground and an erosional unconformity at the base of the overlying car9bonates in the Guandacol area. This unconformity is matched by a type-1 sequence boundary separating the E1 Estero Member from the Soldano Member of the La Laja Formation. The rocks beneath the sequence boundary are white quartz arenites and black shales of a shallow depositional environment. Above the unconformity and above the sequence boundary, two trilobite zones seem to be absent from the sediments of the AP (Palmer pers. comm.). These data indicate that the corresponding erosional event is approximately equivalent in timing to the Hawke Bay event described from the Appalachian margin of Laurentia (Palmer & James 1980). The subsequent Cambrian history of the AP is characterized by rapid but decreasing subsidence until the Dresbachian, and a carbonate factory barely able to keep pace with subsidence (catchup system of Kendall & Schlager 1981). Only during the Late Cambrian is there a keep-up pattern in the sediments. The absence of terrigenous detritus from the Dresbachian onward indicates that the source area had vanished, most probably by onlap of the Upper Cambrian peritidal deposits. Faunal data (Benedetto et al. 1995; Vaccari 1994) indicate that during the Cambrian the AP was faunistically indistinguishable from Laurentia.

Supersequence B (Ibexian-Early Whiterockian) Supersequence B comprises the Lower Ordovician to basal Middle Ordovician carbonate platform sediments attributed to the La Silla Formation and the San Juan Formation. The most striking feature of this supersequence is the total absence of any terrigenous material other than clay.

La Silla Formation (400 m) Lithostratigraphy. The wackestone association is composed of mudstones and wackestones with a scarce and monotonous fauna. Abundant hardgrounds and bedding-parallel trace fossils indicate slow and non-continuous mud accumulation, probably under hypersaline conditions.

42

M. KELLER ET AL.

A more favourable environment is indicated by wackestones and packstones with sponges, receptaculitids, and abundant nautiloids and gastropods. In general, sediments of this association were deposited under low-energy, subtidal conditions. Rocks of the peloidal grainstone association are the most abundant in the La Silla Formation. There are frequent intercalations of intraclast grainstones or intraclast-peloidal grainstones. Most of the intraclasts are of storm origin, the clasts represent the entire spectrum of the microfacies observed in the La Silla Formation. In the oolite association, most of the sediments show abundant cross bedding. Herringbone cross stratification has been observed only locally. Intraclasts are abundant in some of the beds. Thrombolites and microbial mounds are also part of the association. Many of the mounds are onlapped by ooid sands. The presence of herringbone cross-stratification was taken as evidence of a tidal bar origin of the corresponding oolites. Most of the oolites, however, originated as marine sands deposited following major storm events. Storm reworking of semi-lithified sediment is also responsible for the abundant intraclasts found in some of the oolite beds. Microbial laminites and mudstones with abundant mudcracks, pseudomorphs after evaporites, and bird's-eye structures are the most important rocks in the microbial boundstone association. The sediments were formed in higher intertidal and supratidal environments. Intercalated flat pebble breccias are the product of desiccation and local reworking of carbonate layers. Sediments of the La Silla Formation were deposited on a vast platform dominated by shallow subtidal, often restricted environments. In addition, intertidal and rare supratidal areas were present. Although oolites are found throughout the succession, only very few of them can be attributed to shoal systems. The majority of the oolites form sheet-like deposits, typical of redistribution during major storms. Intertidal and supratidal environments are dominated by microbial boundstones and thick storm-induced mudlayers with desiccation cracks and evaporites. Cafias (1995a, b) interpreted the deposits of the La Silla Formation as representing a rimmed shelf. In this model, the oolites are margin-related accumulations. However, there are no indications of a nearby shelf break, nor are the oolites restricted to a single facies belt but are found all across the platform. In the westernmost exposures of the La Silla Formation, intertidal sediments are absent and muddy subtidal rocks prevail. This indicates a ramp like configuration of the platform during

the Lower Ibexian which gently dipped towards the west.

Sequence stratigraphy. Within the La Silla Formation three major sequences are developed (sequences 9-11; Fig. 4). They start with subtidal limestones, mainly of the wackestone association. Up section, packstones and peloidal grainstones are developed. The lower two sequences are topped by microbial laminites and rare thrombolites, whereas the uppermost sequence shows bird's-eye mudstones and thrombolites in this interval. At the base of the lowermost succession, silty dolomites, large intraclasts and extraclasts testify to exposure and erosion of underlying strata. Consequently, this boundary is a type-1 sequence boundary. The nature of the subsequent boundaries is less obvious. Erosional features are very rare and there is no obvious basinward shift in facies, hence they are interpreted to represent type-2 boundaries (Keller 1997). Based on their average duration of 3 - 4 M a , all of these sequences are thirdorder sequences. The lowermost and the uppermost successions are relatively thin and contain abundant intertidal deposits. In contrast, in the middle succession, which is relatively thick, subtidal lithologies are most abundant. This tripartite subdivision of the La Silla Formation has a close match in the coeval strata of the northern Appalachian margin. There, this interval, corresponding to the Lower Ibexian, consists of a predominantly subtidal unit sandwiched between peritidal carbonates (Knight et al. 1995). Similarly, the Chepultepec interval of the southern Appalachians shows a succession of peritidal carbonates, subtidal rocks and, again, peritidal strata (Bova & Read 1987). San Juan Formation (330 m) Lithostratigraphy. The most complex facies association is the reef and reef mound association. It consists of biohermal and biostromal accumulations of sponges, receptaculitids, early stromatoporoids, and algae. Non-reef rocks are rudstones, grainstones, wackestones, and in places bird's-eye mudstones. The rocks of this association have been recently described by Cafias & Carrera (1993), Cafias (1995a, b), Cafias & Keller (1993), Keller & Bordonaro (1993), and Keller & Flfigel (1996). The reefs are concentrated in two horizons, one near the base of the formation and one near the top. The packstone-grainstone association is a lateral equivalent of the reef and reef mound

PRECORDILLERA STRATIGRAPHY association. A few isolated biostromes have been found within this association. Packstones and grainstones contain broken and abraded fragments of a varied fauna. Intraclasts are also present, but are only locally abundant. In a few sections cross-bedded, almost monomict pelmatozoan grainstones have been observed. The wackestone-intraclast packstone association and the wackestone-oncolite association both represent the background sedimentation (wackestones and a few mudstones) on the carbonate platform during the Late Ibexian and Early Whiterockian. The sediments and their organic content point to quiet subtidal conditions still within the photic zone. The intraclast packstones and the oncolite beds are the results of storm events affecting different source areas. Tempestites are rare in the sections of the Sierra de Villicum and Sierra Chica de Zonda, but they are very abundant in the J~chal area. There, the beds originating from tempestites are thick and often show an erosional base, grading, and cross bedding. Farther towards the west beds tend to become thinner, erosional features disappear, and oncoids and intraclasts become less abundant. The nodular-wackestone association is dominated by whole-fossil wackestones and few mudstones; tempestites are absent. The association was deposited below storm wave base, but still within the photic zone as indicated by the abundant and diverse fauna. Similarly, the nodularpackstone association records a relatively deep environment; a few graded beds and shallow water biota in these beds, however, indicate some storm activity. Shale partings, thin shale beds, and the total homogeneization of many beds by bioturbation point to reduced sedimentation rates. The mudstone-wackestone-shale association is characterized by dark colours and platy, thinly bedded rocks. The fauna consists of deepwater trilobites and conodonts. The carbonates are typical deep-water, hemipelagic limestones (Wilson 1969). An additional association was described by Cafias (1995a, b) from the Guandacol area. There, rocks similar to the hemipelagic limestones described above alternate with fine- to medium-bedded, poorly sorted packstones and grainstones. The sedimentary succession is interpreted as a deep-water setting, but still within the reach of tempestites. The sediments of the San Juan Formation were deposited on a carbonate ramp dipping gently towards the west. The easternmost outcrops contain the high-energy shoal-water complex which hosts the main reef accumulations.

43

The upper reef horizon can be traced westward into slightly deeper packstones and grainstones with abundant reef-derived material. A similar pattern holds for all sedimentological units described from the San Juan Formation (Cafias 1995a, b; Keller & Flfigel 1996; Keller 1997). The eastern sections always represent a (slightly) more shallow environment than sections farther to the west. Remarkable, however, is the absence of lagoonal, intertidal, and supratidal deposits, indicating that an important part of the platform is not preserved.

Sequence stratigraphy. The San Juan Formation is composed of two major sequences. The lower one starts above the boundary separating the San Juan Formation from the La Silla Formation. Above this type-2 sequence boundary, several small-scale shallowing-upward cycles form a progradational parasequence set, interpreted as a shelf-margin systems tract. Above the overlying transgressive surface, the lower reef mound horizon formed during a rapid relative rise in sea level. Within this transgressive systems tract, several additional flooding surfaces have been identified. The accompanying deepening finally led to the deposition of a characteristic nodular limestones association deposited on the deep ramp. The most marked deepening is observed in the O. evae conodont zone and was caused by a eustatic sea-level rise of global dimensions (Fortey 1984). The top of this lower nodular limestone association is regarded as the zone of maximum flooding, which marks a change from retrogradation to aggradation and the onset of the highstand systems tract. The next sequence boundary is also well defined by the sudden change from mid-ramp deposits to shallow-water grainstone and packstones. In the La Silla section, a hiatus is presumed to exist between the two facies associations. In addition, the Tremadocian conodont Cordylodus cf. angulatus was found just above the sequence boundary, indicating considerable erosion in the more interior parts of the ramp. Again, the transgressive systems tract hosts the main reef accumulation of the upper reef interval. Several flooding surfaces are present above the reef mounds which indicate successive drowning of the carbonate platform, culminating during the Whiterockian. The drowning succession is characterized by black shales and deep-water limestones, some 10m thick. This succession is overlain by graptolitic black shales. The drowning itself is not a uniform process: drowning of the outer part of the ramp (Guandacol area) was earlier than in the other areas and was coeval with the upper sequence boundary in

44

M. KELLER E T AL.

the San Juan Formation south of Jitchal. Another exception is the Las Chacritas section, where some 60 m of spiculitic wackestones rest on the San Juan Formation and form the drowning succession. Finally, in the Las Aguaditas section a carbonate slope and basin developed above the carbonate platform (Keller et al. 1993a). The very different timing of drowning and the different successions themselves are vestiges of the breakdown of the carbonate platform under the onset of crustal extension. Evolution of supersequence B

In comparison to supersequence A, which culminated with the deposition of a laterally extensive tidal-flat complex, supersequence B documents the return to more open-marine conditions. Third-order sequences tend to become thinner towards the Whiterockian and the character changes from strongly progradational during the latest Cambrian to aggradational towards the top of the San Juan Formation. The sequence boundaries themselves also show an evolutionary trend: type-1 boundaries are restricted to the Late Cambrian, whereas type-2 boundaries characterize the Ibexian and Early Whiterockian. In the La Silla Formation, the sequences are topped by microbial laminites and inter- to supratidal facies. However, the thickness of these late highstand deposits progressively becomes thinner and in the San Juan Formation no intertidal or supratidal facies are found at the top of the sequences. This overall trend in the evolution of the sequences and the corresponding boundaries is accompanied by the change from mostly dolomitic rocks of the Late Cambrian to exclusively limestones during the Early Whiterockian. The carbonate rocks of supersequence B reflect various attempts at drowning the platform. Although the carbonate factory was prolific well into the Whiterockian, the sum of all external factors was finally too strong to permit further carbonate production. The Ibexian La Silla Formation shows a rather uniform facies development. Two important events accompanied the deposition of this unit. Near San Rafael (province of Mendoza), Grenvillian basement attributed to the Cuyania terrane (Ramos 1995) was flooded and a carbonate succession similar to the La Silla Formation and basal San Juan Formation was deposited. The cratonic character of this part of the terrane (Fig. 1) is documented by 80m of limestones which correspond to almost 400 m in the AP proper (Bordonaro et al. 1996). The other important event was the incipient evolution of a slope environment to the west of the

platform, although rocks of this depositional setting are only preserved within the Los Tuneles mega-olistolith.

General aspects of supersequences A and B During the formation of supersequences A and B (Cambrian to basal Middle Ordovician), sedimentation in the AP was controlled by the same factors as those responsible for the evolution of the Appalachian margin. These include a thermally subsiding crust, localized tectonic events (e.g., Hawke Bay event), and the increasing importance of eustasy, which exerted major control on the formation of 13 third-order sequences in the AP (Fig. 4). The corresponding qualitative sea-level curve and its correlation with the curve for the southern Appalachians (Read 1989) is shown in Fig. 5. In addition, the benthic faunas of the whole of supersequence A and most of supersequence B (Cambrian through Late Ibexian) cannot be distinguished from the faunal succession in Laurentia (Benedetto et al. 1995; Fig. 6a-c), but are totally different from those of the terranes today adjacent to the AP. The first non-Laurentian fossils are associated with sponge-algal mounds in the lower San Juan Formation (Vaccari 1995). However, the faunal record and the datasets from Laurentia and the AP are not of the same quality. In the AP, there are almost no data on benthic faunas for most of the Early Ibexian. The Laurentian-type conodont succession of the AP (Lehnert 1997; Lehnert et al. 1997) may help to fill this gap. The compilation of the palaeobiogeographical affinities of benthic macrofossil assemblages of the AP (Benedetto et al. 1995; Fig. 6a) includes data on several groups. These data are not always of uniform quality, because some groups have been recorded only from a short interval and/or are represented only by a few taxa (e.g., sponges: Late Ibexian-Early Whiterockian). In addition, there are no data on brachiopods, bryozoans, pelecypods and sponges before the Late Ibexian. Finally, the individual groups may not be of comparable palaeogeographical sensitivity.

Supersequence C (Mid- and Late Ordovician) Sediments of this supersequence were deposited in two different settings. In the western AP, the developing basin accommodated thick successions of a continental margin and adjacent basin

PRECORDILLERA STRATIGRAPHY

45

Fig. 5. Sea-level curve for the Cambro-Ordovician carbonate platform deposits and its correlation with the Appalachians. The Appalachian sea-level curve is modified from Read (1989). A and B show alternative correlation of earliest Ordovician sequence boundaries in the Argentine Precordillera. environment. In the east, this supersequence developed above the former carbonate platform.

The western basin

The most impressive sediments are those of the Los Sombreros Formation. Black graptolitic shales host various types of mass-flow deposits. Debris-flow deposits, turbidite sandstones and greywackes, mega-breccias, and conglomerates

are volumetrically abundant. However, most spectacular are giant olistoliths, which in places are more than 1 km long and more than 300 m thick. It is in this facies that the Middle and Upper Cambrian deep-water limestone olistoliths are present. Other important components of the mass-flow deposits are clasts of metamorphic basement (Banchig et al. 1990). These clasts are interpreted as having been derived locally (Banchig et al. 1990; Keller 1995). Since in some sections basement clasts are found

46

M. KELLER E T A L .

Fig. 6. (a) Changes in the composition of benthic faunas in the Argentine Precordillera on the genus level (modified from Benedetto et al. 1995). During the Cambrian, the fauna is Laurentian. With the onset of extension during the Whiterockian, the first Avalonian and Baltic elements are observed. Note that an influence of Gondwanan elements cannot be observed before Early Caradoc. In addition, from a faunistic point of view, the Caradoc is the time of maximum isolation indicated by a high percentage of endemic faunas and pandemic genera. (b) The trends in palaeobiogeographical affinities are much better visible on a sketch excluding pandemic faunal elements; (e) includes the aspect that there are no reports on other than Laurentian faunas before the Late Ibexian. Faunal composition contradicts any model which claims for a mid-Ordovician collision of the Argentine Precordillera with Gondwana.

together with clasts of the entire carbonate platform succession, their presence implies the formation of a fault scarp more than 2100m high (the cumulative thickness of the carbonate platform succession). Siliciclastic clasts that might have been derived from a rift-related cover of the basement are only present is the Los Tuneles section, where big blocks of quartzpebble conglomerates are exposed. These blocks are lithologically comparable with sediments of the Ocoee Group, rift-related deposits of the Appalachian margin (R. Hatcher pers. comm.). The scarceness of rift-related clasts in most sections indicates that the basement of the AP was mostly covered by carbonate platform sediments (Keller 1995). All mass-flow deposits are proximal, and the giant olistoliths must have been transported by rockfall. In addition, the sediments of the Los Sombreros Formation trace the former margin of the carbonate platform. However, as the siliciclastic succession is younger (latest Llanvirn-Caradoc; see discussion in Keller 1997), there is no sedimentological continuity between the former platform and the newly developed basin. The depositional environment of the Los Sombreros Formation is characterized by local steep scarps which delivered the olistoliths and many of the conglomerates, and a basin area (the continental rise) where these sediments accumulated. Farther to the west, there are several siliciclastic units deposited in a deep basin. The Portezuelo del Tontal Formation contains abundant conglomerates and turbidites (Spalletti et al. 1989) and is interpreted as a deposit of the continental margin. The Don Polo Formation is a succession of proximal to distal turbidites (Fig. 7) and includes some horizons interpreted as contourites (Astini 1991; Keller 1997). The Alcaparrosa Formation consists of graptolitic and aluminiferous black shales. In addition, it contains basaltic pillow lavas (Fig. 7) with MORB-like characteristics (Haller & Ramos 1984; Kay et al. 1984). Most of these sediments are poorly dated, but they all fall in the range from latest Llanvirn to Caradoc. Although the structural relations between these units are not always clear, the spatial distribution of the sediments exhibits an overall fining towards the west (Fig. 7), culminating with basin shales and pillow basalts. West-directed transport has been determined for the proximal deposits and a SWto S-directed transport for the more distal sediments (Spalletti et al. 1989; Astini 199l). All sedimentological data indicate that the sedimentary succession of the western basin is a rift-drift related succession marking the

PRECORDILLERA STRATIGRAPHY

47

Fig. 7. Tectono-sedimentary environments, evolution of the western basin, and sediment distribution in the Argentine Precordillera following continental break-up. Note that there are places where the rift-onset and the break-up unconformities merge. See text for discussion. (Not to scale)

formation of an Atlantic-type continental margin in the AP during the Mid- and early Late Ordovician. This is also evident from the very high sediment accumulation rates (Fig. 8).

The former platform area Above the drowning unconformity of the carbonate platform, most of the deposits are graptolitic black shales. Only locally did carbonate sedimentation continue (Las Chacritas and Las Aguaditas sections) and even in these successions there is a clear tendency towards deep-water limestones. In most sections then, there is a hiatus or an erosional unconformity above which mass-flow deposits dominate. The corresponding event affected the platform area during the earliest Llandeilo. In the Don Braulio section, three fining-upward successions were deposited (La Cantera Formation, c. 150m). The basal succession starts with massive conglomerates locally containing large slabs of the underlying black shales; up-section, turbidites are found. The upper two successions start with coarse sandstones and arkoses, which pass into shales. Another hiatus separates the La Cantera Formation from the Don Braulio Formation, which contains vestiges of the Ashgillian Gondwana glaciation (Peralta & Carter 1990; Buggisch & Astini 1993).

In the Guandacol area, a similar succession is preserved, but with a thickness of more than 1200 m (Trapiche Group). In the basal part of this succession, limestone slabs up to 70 m across are present (Astini 1991). In addition, several high-angle erosional unconformities have been observed in one of these sections (yon Gosen pers. comm.). All these features point to rapid local uplift and the formation of oversteepened fault scarps. The highly varied sediment accumulation rates above the former platform are documented in Fig. 8. A unique succession is preserved at Rio Sassito where about 40m of mainly Caradoc sediments rest on top of deeply eroded San Juan limestones. The lower part of the succession is composed of about 20 m of siliciclastic rocks, the upper 20m are temperate-water limestones. These limestones apparently formed in a horst position because the adjacent areas are characterized by deep-water siliciclastic material with mass-flow deposits. This outcrops demonstrates that a major period of erosion affected the former platform area after the deposition of the Middle Ordovician drowning succession but prior to the Caradoc. Everywhere in the former platform area, Ordovician rocks of supersequence C are separated from the overlying supersequence D by a major unconformity. In many places the basal bed of the younger succession is marked

48

M. KELLER ET AL.

Fig. 8. Plot of cumulative sediment thicknesses in different areas and settings in the Argentine Precordillera. The most marked differences between the platform area and the western basin occurred during Mid- and Late Ordovician time, caused by rapid subsidence of the newly forming basin. Inlay shows history of thermal subsidence. This plot includes only data from the carbonate platform and, during the Late Ordovician to Devonian, from the former platform area (from Gonzfilez Bonorino & Gonz/Llez Bonorino 1991).

by a chert-pebble conglomerate. The onlap of the younger succession onto the eroded platform is strongly diachronous (Astini & Maretto 1996). The overlying strata, however, show a rather uniform development.

Evolution of supersequence C Supersequence C is bounded below and above by two major unconformities. The basal unconformity marks the end of carbonate platform deposition and the onset of a depositional regime which favoured mass-flow sedimentation with highly variable thicknesses; in places, carbonates were deposited. Local unconformities

are common in all successions. The basal unconformity is approximately coeval with the formation of a continental margin and the adjacent western basin. In our interpretation, supersequence C represents major crustal extension related to rifting and the concomitant formation of horst and graben structures, the evolution of a continental rise and an adjacent basin into which MORB-like basalts were extruded. The major fault scarp along which the Los Sombreros Formation was deposited is evidence of the rifting: a scarp more than 2100m high is incompatible with a compressive regime and must be related to large-scale crustal extension. An extensional regime is corroborated by the

PRECORDILLERA STRATIGRAPHY pillow basalts and the formation of Caradoc carbonates on an isolated horst. Various studies have attributed individual sedimentary successions to crustal extension (Spalletti et al. 1989; Loske 1992; von Gosen 1992; Keller 1995). Similarly, the pillow basalts were interpreted as having formed in an extensional regime (Haller & Ramos 1984; Kay et al. 1984). Dalla Salda et al. (1992b) were the first to interpret the Middle and Upper Ordovician strata as representing rifting and in particular the separation of the AP from Laurentia. A similar argument has been used by Keller & Dickerson (1996). Increasing distance from Laurentia is also apparent in the composition of the faunas (Fig. 6). Avalonian, Baltic, and endemic faunas became increasingly abundant during MidOrdovician times. Laurentian faunas were able to settle on the disappearing terrane until the Late Ordovician. In contrast, the first Gondwanan (Mediterranean) brachiopods did not arrive before the Caradoc, which was the time of maximum faunal isolation (Fig. 6c). Falvey (1974) proposed a comprehensive model for rifting, continental break-up, and the evolution of passive continental margins. This model includes two major unconformities: a riftonset unconformity and a break-up unconformity (Fig. 7). Thermal expansion of the crust and concomitant erosion during initial crustal stretching are responsible for the rift-onset unconformity. In a subsequent stage, crustal expansion leads to the formation of horst and graben structures (rift valley stage), whereas the break-up unconformity marks the transition from rifting to drifting. Consequently, the breakup unconformity is almost coeval with the onset of formation of oceanic crust. In the AP, the major unconformity just above the drowning succession, which marks the base of supersequence C, is here interpreted as the rift-onset unconformity (Fig. 7). In the former platform area, the concomitant erosion cut down to highly varied stratigraphical levels in the San Juan Formation (Fig. 9). In the western basin, this unconformity is correlative with the basal strata of the Los Sombreros Formation, which in turn was deposited during the evolution of a continental slope and rise. The rift-valley stage is well expressed on the former platform where highly varied successions of conglomerates, mass-flow sediments, and sandstones were deposited in grabens or half-grabens. Locally, a carbonate slope developed. On corresponding horsts, successions like that at Rio Sassito were laid down. A widespread hiatus is present in the Caradoc, especially on the former platform. This hiatus might have been caused by lateral heat-

49

flow from the active volcanic centres in the basin. This heat-flow caused uplift and erosion on the platform and the resulting unconformity has to be regarded as the break-up unconformity (Fig. 7). The widespread latest Ashgillian/ Silurian chert-pebble conglomerate at the base of supersequence D marks the break-up unconformity on much of the former platform. However, there are successions (Don Braulio Formation, Empozada Formation sensu Keller 1997) where there are relatively thick Ashgillian sediments preserved just above the break-up unconformity.

Supersequence D Silurian deposits are widespread between Jfichal in the north and the Rio San Juan in the south. They are absent in the Guandacol area and seem to be absent in the Mendoza area. Two different environments are evident. Along the eastern margin of the AP (Sierra Chica de Zonda and Sierra de Villicum), a deep graben accommodated mass-flow deposits (Rinconada Formation). Clasts and huge olistoliths mirror the underlying stratigraphy and document erosion down to the San Juan limestones. Consequently, the succession of clasts and olistoliths shows an inverse stratigraphy. In addition, there are abundant clasts with lithologies not found in the platform succession (Loske 1992; yon Gosen et al. 1995), mainly acidic to intermediate magmatic rocks. The thickness of the sediments (conservative estimates are on the order of 1000m, von Gosen et al. 1995), abundant slump and slide structures, and the huge olistoliths document rapid subsidence of a basin and uplift of the adjacent shoulders. The other sedimentary environment is a vast siliciclastic platform with facies spanning the shoreface to areas well below storm wave base (Acefiolaza & Peralta 1986; Peralta 1990; Peralta & Carter 1990; Astini & Maretto 1996). Well into the Devonian, this platform had a clear N-S polarity. Sediments are thickest around Jfichal and get thinner towards the Rio San Juan. In addition, there is an increasing number of hiatuses towards the south, and the time gap represented by the hiatuses increases in the same direction (Benedetto et al. 1992; Herrera 1993; Peralta 1993). The J~chal area was obviously a depocentre, whereas the Rio San Juan area was a structural high. Furque & Caball6 (1990) described N-Strending ridges which subdivided the siliciclastic platform during the Silurian. Although this pattern is indeed visible in several transects

50

M. KELLER E T AL.

from west to east across the platform, the N-S polarity is much more pronounced. The Devonian sediments document a slow transition from siliciclastic shelf sedimentation to the deposition of turbidites. The Lower Devonian Talacasto Formation is predominantly characterized by shales in its lower part; up-section, sandstones dominate. The overlying Punta Negra Formation consists of flysch-like deposits (Gonzfilez Bonorino 1975) attributed to a fan-delta complex. The contact between the Talacasto Formation and the Punta Negra Formation is strongly diachronous (Herrera 1993), although it is not clear whether this is true diachronism or whether it reflects a major episode of erosion prior to the onset of flysch formation. The loss of section from the top down in a southward direction might favour the erosional interpretation whereas the onset of flysch deposition during the Early Devonian in the Mendoza area might be taken to indicate a northward migration of the fan deltas and hence a diachronous evolution. Within supersequence D, there is an interesting evolution in the heavy mineral spectra (Loske 1992). The Silurian (and Ordovician) sediments are dominated by a low-diversity and low-quantity association of weathering-resistant minerals, such as zircons and tourmalines. In contrast, the Devonian clastic rocks received a high-diversity and high-quantity association, with garnet, zoisite, and apatite, among others. Following Loske (1992) this change mirrors the change in the source rocks from a mature hinterland with unmetamorphosed or slightly metamorphosed rocks, to a hinterland where magmatic and metamorphic rocks were exposed. Zircon ages from the Devonian sediments (1.1 Ga, Loske 1995) show that this magmatic/ metamorphic hinterland might have been the basement of the AP itself (the Cuyania terrane).

during the Wenlockian, the AP becomes faunistically indistinguishable from Gondwana (see Benedetto this volume). The Devonian sediments indicate a change from platform sedimentation to the evolution of a flysch trough which received detritus from a rising metamorphic or magmatic hinterland. Palaeocurrent data (Gonzfilez Bonorino 1975) suggest northeastern and southeastern source areas which fed two important fan deltas. Supersequence D also saw the onset of deformation in the AP, marked by a thermal event which affected the western margin of the AP between 425 Ma and 410 Ma (Buggisch et al. 1994). Conodont CAI data (Keller et al. 1993b; Fig. 9) show an increase from 150~ at the eastern margin of the AP to more than 300~ in the continental rise deposits. The concomitant compressional deformation caused west-vergent structures in strata as young as the Devonian Talacasto Formation (von Gosen 1995). Westdirected thrusts and imbrications are also present at the eastern margin of the AP (von Gosen 1992; von Gosen et al. 1995). These structures, which are also visible in seismic lines (Cominguez & Ramos 1990), have a post-Silurian but pre-Late Carboniferous age.

Evolution of supersequence D

Loske (1992) and von Gosen et al. (1995) attributed the sediments of the Rinconada Formation to a depocentre formed by crustal extension. During the Silurian, this extension obviously also affected the remainder of the platform area. The basins and ridges described by Furque & Caball6 (1990) might reflect the formation of tilt blocks striking N-S. The asymmetry between the Jfichal area and the structural high along the Rio San Juan, however, documents an overall tilting of the shelf in a N-S direction. Faunistically, the mid- to Late Silurian marks the change towards entirely Gondwanan elements. With the advent of the Clarkeia fauna

Fig. 9. 'Metamorphism' and level of erosion on the carbonate platform due to block faulting. CAI, conodont colour alteration index. Thermal alteration is weak at the eastern border of the Argentine Precordillera and increases towards the west (locations as in Fig. 2).

PRECORDILLERA STRATIGRAPHY

The plate-tectonic background of sedimentation in the Argentine PrecordUlera It has become widely accepted that the AP is a Laurentia-derived terrane (Keppie 1991; Astini et al. 1995; Dalziel & Dalla Salda 1996; Dalziel et al. 1996; Keller & Dickerson 1996; Thomas 1996) and that it originated in the Ouachita embayment as originally proposed by Dalla Salda et al. (1992a, b). However, the timing of separation of the AP and the plate-tectonic scenario involved are highly controversial (Dalla Salda 1992a, b; Dalziel et al. 1994; Astini et al. 1995; Keller & Dickerson 1996; Dalziel 1997; Keller 1997). The rift-drift transition along the Appalachian margin is supposed to have taken place around the Precambrian-Cambrian boundary (Bond et al. 1984, 1989; Read 1989; Thomas 1996). In contrast, rifting along the Ouachita margin is younger (Thomas 1996) and a passive margin was not established before the MidCambrian. Intracontinental graben systems indicative of incipient rifting are an important feature along the Appalachian-Ouchita margin (Thomas 1991). Typical sediments are finegrained continental redbeds and local evaporites. Sedimentological evidence from the AP records the following steps. (1) Crustal thinning during the Early Cambrian, indicated by the redbeds and evaporites of the Cerro Totora Formation. From the late B o n n i a - O l e n e l l u s chron onwards, carbonate sedimentation prevailed. (2) The occurrence of a widespread hiatus in the AP possibly equivalent to the Hawke Bay regression event and hence confirming its position on Laurentian crust at that time. (3) During the Glossopleura chron of the Mid-Cambrian, a fully developed carbonate factory became established. We interpret the Early/Mid-Cambrian history of the AP as a record of the formation of an intracontinental rift which separated the AP to some extent from the Ouachita margin. (4) The newly formed crustal fragment existed as a Laurentian marginal plateau (Lister et al. 1991). Marginal plateaux are large, relatively unstructured crustal fragments, which owe their characteristics to a mid-crustal detachment and the concomitant pull-out of middle and lower crust from beneath the future marginal plateau (Lister et al. 1991). These plateaux are often separated from the main continent by a ramp syncline or a hanging wall basin, which may be relatively deep. In the AP, carbonate sedimentation started on the newly developed marginal plateau, whereas the deep-

51

water limestones were deposited in the intervening hanging-wall basin. (5) During the Late Dresbachian, much earlier than in other parts of Laurentia, siliciclastic input from a crystalline source area had ceased, testifying to flooding of the plateau. (6) From the Dresbachian to the Early Whiterock9 sedimentation on the carbonate platform proper was almost exclusively controlled by eustasy. The marginal-plateau model helps to explain the absence of margin-derived megabreccias and other mass-flow deposits. The gradient between the platform and the slowly subsiding basin to the west was not great enough to permit gravity-driven processes to operate on a large scale. Supersequence A records the initial intracontinental rifting, the formation of the marginal plateau accommodating the carbonate platform, and the evolution of a hanging wall basin in which deep-water limestones were deposited (Fig. 10). Sedimentology and faunal associations cannot be distinguished from mainland Laurentia. In its higher part, the supersequence documents the loss of a source for terrigenous input and the evolution of a progradational carbonate complex which covered the entire platform. The Early and earliest Mid-Ordovician saw the following events.

my 360 370

~

380

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390

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400

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Punta Negra Flysch > i0J ~O

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410 thermal overprint

,20

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g- ~ mid

chertpebbleconglQmerate

Villicurn graben

9Q 9 I9

440

_ _

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470

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~

VVV~VV ~ pillow "basalts

rift onset u ncon f

rifting

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~

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Fig. 10. Tectonosedimentary evolution of the Argentine Precordillera as discussed in this paper.

i

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52

M. KELLER E T A L .

(7) Successive return to an aggradational carbonate system which escaped drowning several times. The corresponding increase in relative sea level is probably a combination of lithospheric cooling of the marginal plateau and of eustasy, as many of the Laurentian sea-level events are recognized in the AP (Keller 1997). (8) The carbonate platform was drowned during the Whiterockian. Sea-level rise was an important factor, but facies successions in combination with sequence stratigraphy demonstrate that there was also a tectonic component reflecting crustal extension. Supersequence B as a whole reflects the demise of the carbonate platform until its complete drowning. During deposition of supersequences A and B, the history of thermal subsidence and of sediment accumulation (Figs 8 & 10) is characteristic of passive-margin evolution. The Mid- and Late Ordovician deposits mirror a climax of tectonic activity. (9) The onset of crustal extension occurred during the earliest Mid-Ordovician. (10) The climax of rifting occurred during the late Mid- and the earliest Late Ordovician. It resulted in major block movements in the former carbonate platform area and the evolution of an Atlantic-type continental margin with fault scarps locally more than 2100 m high. The pillow basalts in the western basin are interpreted as remnants of the adjacent ocean floor. (11) During the Ashgillian, after its separation from Laurentia, the AP occupied a palaeogeographical position in higher latitudes within the reach of icebergs, and hence ice-rafted sediments. According to Brenchley & Newall (1984), icebergs may have drifted as far north as 30~ during the Late Ordovician glaciation. Consequently, ice-rafted sediments alone (such as the diamictites of the Don Braulio Formation) cannot be taken as evidence the accretion of the AP to Gondwana. Faunal data and the change of the palaeogeographical affinities of the benthic associations display strikingly the timing of separation and the subsequent drift towards isolation of the AP (Fig. 6a). This is even more evident if pandemic faunal elements are excluded from the diagrams (Fig. 6b-c). Supersequence C reveals the continental break-up and the final separation of the AP from Laurentia and the complete opening of the Ouachita basin (see Dickerson & Keller this volume). The supersequence is bounded at the base by the rift-onset unconformity and at the top by the break-up unconformity (Fig. 10). The contrasting sediment accumulation rates

in the former platform area and the newly developing basin are documented in Fig. 8. The Silurian and Devonian history include the following. (12) The drift to high latitudes and the approach to Gondwana. This is shown by the Mid-Late Silurian Clarkeia fauna, because this fauna is only present in a few areas of highlatitude Gondwana (Cocks & Fortey 1990). (13) There is a marked change from siliciclastic platform deposition to flysch sedimentation during the Devonian. (14) Heavy-mineral populations in the Lower Palaeozoic strata indicate that a predominantly sedimentary hinterland was recycled prior to the (Mid-) Devonian. Subsequently, however, a source area was present which was composed of granitoids and metamorphic rocks (Loske 1992; Kury 1993; Astini & Maretto 1996). This change of source rock exposure is here interpreted as an effect of the final accretion of the AP to Gondwana. (15) The Late Silurian-Early Devonian metamorphism and the corresponding deformation in the western part of the AP indicate the onset of structural deformation (Fig. 10). They are interpreted as reflecting closure of the western ocean and collision with the Chilenia terrane (Ramos et al. 1984, 1986), whereas concomitant structures in the east may be related to the accretion of the AP to Gondwana (von Gosen 1992). This scenario is supported by the magma evolution and igneous activity in the Sierras Pampeanas indicating continuous subduction into the Silurian, changing gradually towards collision during Late Silurian-Devonian times (Rapela et al. 1992). It seems logical that the separation of the AP from Laurentia and the formation of oceanic crust must have been accompanied by subduction of the opposite side of the terrane. This in turn may have been subduction beneath the Sierras Pampeanas, causing the corresponding magmatism there (see Saavedra et al. this volume and Pankhurst et al. this volume). (16) The accretion of the AP to Gondwana was terminated prior to the deposition of the Late Carboniferous molasse sediments. Supersequence D is an expression of the final approach and the subsequent accretion of the AP to Gondwana, but also of the approach of Chilenia in the west. The Early Palaeozoic history of the AP resulted from the complex plate interactions between Laurentia, western Gondwana, and the Chilenia terrane (Ramos et al. 1984, 1986; Astini et al. 1996). We are convinced with some certainty that sedimentological evidence and the faunal data from the AP exclude models

PRECORDILLERA STRATIGRAPHY favouring M i d - O r d o v i c i a n collision of the A P terrane (Astini et al. 1995), of the C u y a n i a terrane ( R a m o s 1995), or of the Occidentalia terrane (Dalla Salda et al. 1992a, b) with the eastern Sierras P a m p e a n a s . A scenario as presented by Astini et al. (1995, 1996) and especially its timing is also unlikely. Based on our sedimentological and faunistic data a n d considering the structural evidence (von G o s e n 1992, 1995; von G o s e n et al. 1995; C o m i n g u e z & R a m o s 1990), we favour a m o d e l in which the A P was separated from L a u r e n t i a during the Ordovician and afterwards slowly a p p r o a c h e d G o n d w a n a during the Silurian and D e v o n i a n . This paper is an output of a large project of the University of Erlangen on the evolution of the AP. We thank W. von Gosen and S. Krumm for many helpful discussions and support. We also thank all our Argentinian and non-Argentinian colleagues and friends, who participated in one form or another, for their co-operation and hospitality. The support of project funds to all three of us from the Deutsche Forschungsgemeinschaft is gratefully acknowledged. This paper benefited from reviews by S. Flint (Liverpool) and an anonymous reviewer.

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Early Palaeozoic brachiopods and associated shelly faunas from western Gondwana: their bearing on the geodynamic history of the pre-Andean margin JUAN

L. B E N E D E T T O

Cdtedra de Estratigrafia y Geologia Histdriea, Facultad de Ciencias Exactas, Fisicas y Naturales, Universidad Nacional de C6rdoba, CONICET. Av. VOlez Sarsfield 299, 5000 C6rdoba, Argentina Abstract: The strong provincialism exhibited by early Ordovician to Devonian brachiopod

faunas provides an independent tool for testing palaeogeographical hypotheses. Patterns of biogeographical affinities of early Palaeozoic brachiopods from the Argentine Precordillera, the Sierra de Famatina and the Central Andean basin (NW Argentina, Bolivia, southern Peru) suggest that the pre-Andean margin was linked to northern Iapetus Ocean history. The low-grade Vendian-Early Cambrian Puncoviscana Formation and the broadly coeval rocks from the Carolina and Gander terranes may represent sedimentation in a narrow basin developed between Laurentia and Gondwana during rifting. Palaeontological evidence suggests that the Precordillera developed on the passive margin of Laurentia but moved away as an independent plate during the Ordovician, the hypothesis of the Precordillera as a Laurentian plateau does not explain faunal differences between their Late Arenig-Llanvirn brachiopods. Ordovician subduction beneath Gondwana resulted in formation of the Famatina-Puna-Avalonia volcanic-arc system. Affinities between Celtic brachiopods of this age in volcaniclastic rocks from South America (Famatina) and eastern North America (Maine, Gander, Central Newfoundland) suggest some geographical continuity between them, consistent with palaeomagnetic evidence. It is proposed that accretion of allochthonous terranes to eastern Laurentia was related to collision with the northwestern corner of South America in late Ordovician time, an idea supported by the affinities of Silurian and Devonian brachiopods from Venezuela and Colombia.

The term 'Southern Proto-Atlantic' was first used by Benedetto & Sfinchez (1979) for the Cambro-Ordovician sea-way between the N W corner of Gondwana (western Venezuela and Colombia) and southern Laurentia. This name, and the more widely accepted 'Gondwanan Iapetus' or 'Southern Iapetus' (Dalziel 1992), implicitly admitted the existence of a palaeocontinent to the west of the Andean region of Gondwana (cf. Dalmayrac et al. 1980). In searching for the counterpart of the early Palaeozoic Appalachian rifted margin, Bond et al. (1984) suggested, on the basis of the striking similarities between the Cambrian trilobites of the central Appalachians and western Argentina (Precordillera basin), that the two areas could have been juxtaposed. Similarities between the Cambrian stratigraphy and trilobite faunas of the northern Appalachians and the Argentine Precordillera, and their stratigraphical divergence since the early mid-Silurian, were also noted by Ramos et al. (1986), who interpreted the Precordillera as a displaced terrane sutured with the Pampeanas terrane before the Carboniferous.

A new palaeobiogeographical picture was proposed by Benedetto (1985) and Herrera & Benedetto (1991), using the information from the rich Arenig-Llanvirn faunas of the San Juan Limestone in the Precordillera. The peculiar combination of genera belonging to the Toquima-Table Head, Baltic and Celtic realms exhibited by these brachiopod faunas, led us to propose a model (Fig. la) in which the Andean margin of Gondwana (including the Precordillera basin) is at nearly the same longitude as the southern Appalachians. This configuration (achieved from the classic reconstructions by means of a palaeomagnetically acceptable clockwise rotation of Gondwana of c. 70 ~ explains faunal exchange between the Precordillera and the Laurentian and Baltic-Avalonian margins of the Iapetus Ocean. On the basis of nonbiological evidence, a very similar CambroOrdovician configuration was proposed by Dalziel (1991), who delineated a wander path for Laurentia clockwise around the protoAndean margin from its original hypothetical juxtaposition with the east Antarctica-Australia margin in the Neoproterozoic (Moores 1991).

BENEDETTO, J. L. 1998. Early Palaeozoic brachiopods and associated shelly faunas from western Gondwana: their bearing on the geodynamic history of the pre-Andean margin. In: PANKHURST,R. J. & RAPELA,C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 57-83.

58

J. L. B E N E D E T T O

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Fig. 1. Some palaeogeographical hypotheses mentioned in the text: (a) Early Ordovician reconstruction after Herrera & Benedetto (1991); (b) 'Microcontinental model' (A-B) and 'paired rift margins' model (C-D), from Benedetto (1993); (e) Early Devonian configuration after Benedetto (1984) and distribution of selected brachiopod genera.

BRACHIOPODS AND PRE-ANDEAN MARGIN In biogeographical analysis of the Early Palaeozoic, I pointed out that faunal affinities between Laurentia and the Precordillera are particularly strong in the Cambrian-earliest Ordovician, but show a clear pattern of divergence through the Ordovician (Benedetto 1993). Conversely, non-Laurentian forms (including Celtic, Baltic and Mediterranean brachiopod taxa) progressively increase in number from Late Arenig to the Ashgill. This biogeographical pattern seems to be compatible with only two palaeogeographical scenarios (Fig. lb, from Benedetto 1993), characterized by the following facts: (1) the Precordillera and Laurentia have moved apart since the Cambrian, and (2) the Precordillera represents the conjugate riftmargin to that of the southern Appalachians. In one of these models (Fig. lb, A-B), the Precordillera is interpreted as a far-travelled microplate, rifted off from the southern Appalachians margin in the early Cambrian and colliding with Gondwana by the middle Ordovician. This was the first hypothesis attempting to explain both the changes in provincialism and the transition from a tropical belt in the CambrianEarly Ordovician to a peri-glacial, high latitude zone in the Late Ordovician, suggested by the lithological succession. Finally, this model established a causal relationship between the Early Ordovician Famatinian volcanic belt and the subduction and subsequent accretion of the Precordillera. This geodynamic interpretation, summarized by Benedetto & Astini (1993) and refined and amplified by Astini et al. (1995) and Thomas & Astini (1996), is known as the 'microcontinental model', or more informally the 'funeral ship' scenario (Dalziel 1997). The other biogeographically compatible reconstruction (Fig. l b, C-D) mainly differs in that the Precordillera is interpreted as an autochthonous part of the Gondwana continent, which rifted away as a whole from the Laurentian margin. The 550 Ma configuration is quite similar to the 545 Ma geologically constrained reconstruction proposed by Dalziel (1997, fig. 15), although in the latter the Precordillera is considered as part of the Laurentian 'Texas Plateau'. In this paper Palaeozoic palaeogeographical models of the southern Iapetus ocean are tested against the palaeontological evidence accumulated from the Cambrian to Silurian successions of western Argentina and Bolivia (Precordillera, Famatina Range and Central Andean basin). The goal of this analysis is to determine whether the known fossil record forms a pattern consistent with the reconstructions based on geophysical and/or geological data. Another objective of

59

the work is to try to integrate the evolution of the Precordillera terrane to the early Palaeozoic history of the NW Argentina-Chile-Bolivia segment of the pre-Andean margin and to search for a biogeographically more consistent disposition of volcanic islands bearing Celtic Realm brachiopods.

Brachiopod dispersion and continental separations Brachiopods are benthic organisms with a planktonic larval stage. Recent articulate brachiopods possess larvae with a very short, nonpelagic stage which may last from a few hours to not more than 2 days (Scheltema 1977). Dispersal potential is greater in groups having teleplanic larvae, such as inarticulate brachiopods and several gastropods, bivalves, crustaceans, and echinoderms. In these benthic invertebrates, pelagic development is commonly of 4-6 weeks, so that with an average current velocity of 1 k m h -1, larvae may be transported 670 to 1000 km. The wide geographical distribution of many species of Palaeozoic Brachiopoda strongly suggests that they may have had longdispersal, planktotrophic larvae (Rong et al. 1995); such organisms can be dispersed over long distances along shorelines and throughout epicontinental seas. An important barrier to the colonization of new areas, however, is the variation of water temperature across the latitudinal belts (Scheltema 1977). Dispersion across ocean basins has been demonstrated for some living echinoderms, bivalves and gastropods, which may attain amphi-Atlantic distribution. Such long-dispersal capability is directly related to the duration of larval development and current velocity. However, many widespread benthic invertebrates have a very short planktonic stage or completely lack them. If this was the case for Early Palaeozoic brachiopods, their dispersion could be achieved by rafting on drift materials (e.g., pumice, pyroclastic rocks or seaweed), as has been demonstrated for many groups of marine invertebrates (Jokiel 1990) and was probably important for shallow water organisms in the Ordovician (Neuman 1984). Also important is the presence in the open ocean of 'standing points', such as volcanic islands, which encourage larvae settlement and enlargement of the area occupied by a population of organisms. The possibility of migration from one island to another depends on the sense of oceanic currents, so that understanding the distribution of benthic organisms in the past

60

J. L. BENEDETTO

requires estimation of palaeo-circulation patterns for each time interval.

Palaeobiogeographical evidence Until the early 1980s, brachiopods from the wide-spread Ordovician successions of the Precordillera basin were virtually unknown. Now, more than 60 Ordovician genera have been described, 34 of which come from the ArenigLlanvirn San Juan Limestone, while the remaining taxa are from the overlying mid-late Ordovician mixed carbonate-clastic or terrigenous successions (Las Plantas, Las Aguaditas,

Trapiche, La Cantera and Don Braulio formations; see correlation chart Fig. 2). About 20 genera are known from the early Ordovician volcaniclastic rocks of the Sierra de Famatina and 11 genera from the less-known, thick siliciclastic sequences of NW Argentina and Bolivia. This background now permits a more confident assessment of the palaeontological constraints on southern Iapetus reconstructions. Genera and species recorded from each basin and the corresponding bibliographical source are summarized in Tables 1, 2 and 3, which include some recently discovered, unpublished taxa and re-assignement of previously described forms. The geographical location of early Palaeozoic sedimentary basins and other geological features

Fig. 2. Correlation chart of main early Palaeozoic stratigraphic units of the Precordillera, Famatina and Central Andean basins. Conodont zones from Albanesi et al. (1995a), brachiopod biozones from Herrera & Benedetto (1990).

BRACHIOPODS AND PRE-ANDEAN

MARGIN

61

Table 1. Arenigian and early Llanvirn brachiopods from the Precordillera basin Taxa

Biozones A

Orthidiella sp.* Orthidium geniculatum Herrera & Benedetto Orthidium sp. Hesperonomia sp Notorthis termalis (Herrera & Benedetto)* Nanorthis ornatus Herrera & Benedetto Nanorthisfragilis Benedetto Paralenorthis vulgaris Herrera & Benedetto Paralenorthis sp. Archaeorthis sp. Oligorthis sp.* Productorthis cienagaensis Herrera & Benedetto Platystrophia minuta Benedetto & Herrera Platystrophia sp. Phragmorthis sp. Skenidioides sp.* Paurorthis ellipticus Herrera & Benedetto Pomatotrema mesocosta Benedetto Pomatotrema talacastoensis Benedetto Tritoechia azulensis Benedetto Tritoechia inaequicostata Benedetto Tritoechia sp. Acanthotoechia n.sp.* Ahtiella argentina Benedetto & Herrera Sanjuanella plicata Benedetto & Herrera Inversella (Reinversella) arancibiai Herr. & Bened. Aporthophyla sp.* Taffia anomala Benedetto & Herrera Leptella (Petroria) rugosa Wilson Leptetta (Leptella) variabilis Benedetto & Herrera Leptella (Leptella) alata Benedetto & Herrera Leptella (Leptella) costellata Benedetto & Herrera Leptella (Leptella)plana Benedetto & Herrera Huacoella radiata Benedetto & Herrera Niquivilia extensa Benedetto & Herrera Leptellina spp. Syntrophia sp.* Rugostrophia sp.* Cuparius sp.* Camerella sp.* Idiostrophia cf. I. perfecta Cooper*

H

N

M

Ah x x

x

• x

x x

x

x x x

x

x x

x

x x • x x x

x

x

x x x x x x x

• • x • • • x • •

• •

•

x x

• x x • x • • •

• x •

Biozones: A, Archaeorthis; H, Huacoella; N, Niquivilia; M, Monorthis; Ah, Ahtiella. Data from: Benedetto (1987), Benedetto & Herrera (1986, 1987a, b, 1993) and Herrera & Benedetto (1987, 1988). * Unpublished taxa.

discussed in this p a p e r are s h o w n in Fig. 3. W i t h s o m e m o d i f i c a t i o n s , the E a r l y O r d o v i c i a n biost r a t i g r a p h i c a l f r a m e w o r k (Fig. 2) is t h a t p r o p o s e d b y H e r r e r a & B e n e d e t t o (1991), w h o r e c o g n i z e d five b r a c h i o p o d a s s e m b l a g e zones c a l i b r a t e d b y m e a n s o f the a s s o c i a t e d c o n o d o n t s ( A l b a n e s i et al. 1995a; L e h n e r t 1995 a n d references therein).

P a l a e o b i o g e o g r a p h i c a l evidence discussed in this p a p e r is largely b a s e d o n the studies by H e r r e r a & B e n e d e t t o (1991), B e n e d e t t o et al. (1995), Astini et al. (1995) a n d B e n e d e t t o (1995a, b) c o n c e r n i n g the P r e c o r d i l l e r a n brac h i o p o d s . T h e b i o g e o g r a p h i c a l affinities o f N W A r g e n t i n a a n d F a m a t i n a b r a c h i o p o d s a n d associated shelly f a u n a s w e r e a n a l y s e d b y V a c c a r i

62

J. L. BENEDETTO Table 2. Arenigian brachiopods from the Famatina and Central Andean basin (CAB) Famatina Hesperonomia orientalis Benedetto Hesperonomia sp. Hesperonomiella sp.* Monorthis cf. M. menapiae (Davidson) Monorthis coloradoensis Benedettot Nanorthis spp. Paralenorthis riojanus (Levy & Nullo) Paralenorthis altiplanicus Benedetto Paralenorthis immitatrix Havli~ek & Branisa Incorthis boliviana Havli6ek & Branisa Incorthis aft. L maroccana Mergl Incorthis sp. Famatinorthis turneri Levy & Nullo Desmorthis segnis Havli6ek & Branisa Glyptorthis imbrex Havli~ek & Branisa Glyptorthis cf. G. imbrex Havli6ek and Branisa Pleurorthis ? sp. Ffynnonia sp.* Prantlina ? sp.* Skenidioides sp. Crossiskenidium sp.* Salopia ? lipanensis Benedetto Tritoechia sp. Euorthisina sp. Camerella sp.*

CAB x

• x • • • x • • • • x • • x • • x • • • x x •

x •

* Unpublished taxa. t Probably younger. Data from Levy & Nullo (1973), Havli6ek & Branisa (1980) and Benedetto (1994, in press b).

al. (1993), Benedetto (1994), Benedetto & Sfinchez (1996a) and Benedetto (in press b). A complete summary of the Precordilleran benthic faunas, including sponges, bryozoans, brachiopods, molluscs and trilobites was presented by Benedetto et al. (1995) and a more comprehensive review of provincial affinities of the Precordillera terrane and related areas was recently concluded (Benedetto et al. in press). Provincialism of Cambrian-Ordovician trilobite faunas from Argentina was reviewed by Bordonaro (1992), Vaccari et al. (1993), Waisfeld (1995), Sfinchez & Waisfeld (1995), Vaccari (1994, 1995) and Astini et al. (1995). Brachiopod biogeographical entities employed in this paper are those recognized by Williams (1973). The term 'Toquima-Table Head Realm' (Ross & Ingham 1970) is used here as equivalent to Williams's 'American Realm', the latter being used in a wider sense than that of Neuman & Harper (1992) who proposed it to replace the William's 'NW American Province'. The following discussion reviews the biogeographical links that constrain the positions and et

motions of island arcs, microplates and continents involved in the southern Iapetus palaeogeography. Trilobite and conodont data are considered, especially for the Cambrian Period, in which brachiopods are nearly absent, and in the time-intervals where the brachiopods are scarce or still poorly known. Data are reviewed separately for each geological region in ascending stratigraphical order. Following Fortey et al. (1995) the Llandeilo is considered as the upper stage of the Llanvirn Series. P r e c o r d i l l e r a basin

(1) Early to Late Cambrian trilobite faunas from the Precordillera basin and Laurentia belong to the same province (Borrello 1971). Recent studies by Vaccari (1994, 1995) demonstrate a strong affinity not only at generic level but also at specific level. This is particularly evident in the Dreisbachian faunas (Crepicephalus Zone) of the northern Precordillera, which share several species with the Appalachians (Vaccari 1994, 1995).

BRACHIOPODS AND PRE-ANDEAN MARGIN

63

Table 3. Caradocian brachiopods from the Precordillera (Precord.) and the Central Andean basin (CAB) Precord. Hesperorthis sp. Dinorthis sp. Dinorthis flabellum (Sowerby) Howellites cf. H. macrostoma (Barrande) Drabovia ? sp. Campylorthis gualcamayensis Benedetto Oanduporella alamensis Benedetto Oanduporella sp. Tissintia robusta Benedetto Tissintia canalifera Havli~ek & Branisa* Eorhipidomella cardocanalis Havli6ek & Branisa Destombesium pacochicoense Havli~ek & Branisa Drabovinella curiosa Havli~ek & Branisa Drabovinella cf. D. erratica (Davidson) Drabovinella sp. Hirnantia? sp. Bicuspina deflecta Benedetto Bicu~pina riojana Benedetto Anchoramena cristata Benedetto Anchoramena? sp. S. (Sowerbyella) cf. S. (S.) sericea (J. de C. Sowerby) Aegiromena glacialis Benedetto Aegiromena corolla Havli6ek & Branisa* Anoptambonites sp.t Oepikoides notus Benedetto Bystromena? protegula Benedetto Camerella sp. Rostricellula sp.

CAB

• • x x • x x

x x • • • x x • x •

* Considered as late Llanvirn by Havli6ek & Branisa (1980). t Unpublished. Data from Benedetto (1995a, 1997a, in press a), Su&rez Soruco & Benedetto (1996) and Havli6ek & Branisa (1980).

(2) The unique trilobite species described from the earliest Ordovician strata of the Precordillera (La Silla Formation), Plethopeltis obtusus, is restricted to the Laurentian shallowwater carbonate platform (Vaccari 1994). The overlying early Ibexian (Skullrockian) limestones contain the conodont Clavohamulus hintzei, a typical form of the North American MidContinent Province; outside Laurentia it is only recorded from Georgina basin of Queensland, eastern Australia (Lehnert et al. 1997). (3) The earliest Arenig faunas from the Precordillera (Archaeorthis Zone) include the typical Laurentian brachiopods Syntrophia, Orthidium and Nothorthis. Syntrophia is recorded from nearly coeval limestones of Arkansas, Vermont and Quebec, as well as from the allochthonous Cow Head Group of western Newfoundland (Ross & James 1987) and from the central Newfoundland volcaniclastic rocks (Neuman 1984). Outside North America, Orthidium was

recorded from the Tourmakeady Limestone of Co. Mayo, Ireland (Williams & Curry 1985), which was interpreted as a dismembered part of Laurentia, and from northeastern China in the Altai-Hinggan mobile belt (Chen& Rong 1992). The Chinese assemblage was considered by Neuman & Harper (1992) as closely related to those of the Celtic province. Nothorthis is also present in western Newfoundland, in the Fort Cassin Formation of Vermont (where it is associated with Syntrophia), in the L6vis Shale of Quebec (Cooper 1956), and in the abovementioned Tourmakeady Limestone of Ireland. The Baltic occurrence of Nothorthis in late Arenig (B. naris Zone) rocks of Estonia (Rubel 1961), is the only one known outside Laurentia. (4) Mid-Late Arenig bachiopod assemblages (Huacoella and Niquivilia zones, see Table 1) are characterized by the abundance of Leptellinae, the pentamerid Cuparius (a typical Laurentian form, but probably present in Kazahkstan), and

64

J. L BENEDETTO

Fig. 3. Location map of early Palaeozoic sedimentary basins of Central and Southern Andes and Puna volcanic belt ('Faja eruptiva de la Puna').

the first record of Platvstrophia (a genus diagnostic of the Baltic Province but widespread in beds younger than Llanvirn). Precordilleran leptellinines include Leptella (Leptella), present in Laurentia and Ireland, and the exclusively Laurentian subgenus Leptella (Petroria), which is represented in the Precordillera by P. (P.) rugosa, the same species founded in the British Columbia. Another distinctive feature is the development of an endemic stock of leptellinines represented by the genera Huacoella and Niquivilia. The species Leptella corbetti, described by Laurie (1991) from coeval rocks of Tasmania (0. evae zone), may in fact belong to Niquivilia (Benedetto & Sfinchez 1996a). Another biogeographical link with the Scoto-Appalachian faunas is Acanthotoechia, a genus recorded only from the Irish Tourmakeady Limestone. (5) The dominance of the genus Monorthis and the addition of Oligorthis and Aporthophyla to several forms recorded from the underlying zones are the most distinctive features of the latest Arenig. Monorthis is restricted to the Ogof Hen and Treiorwerth formations of Anglesey (Bates 1968), the type-locality of the Celtic Province. The species Monorthis noctilio, mentioned by Gutierrez-Marco et al. (1984) from the

Llanvirn of Spain probably belongs to Hesperonomia. The true Oligorthis has only been reported from the Arbuckle Limestone of Oklahoma (Cooper 1956), but Oligorthis? sp. was recently discovered in early Arenig limestones of the Canning Basin, Western Australia (Laurie 1997). (6) As Herrera & Benedetto (1991) pointed out, the uppermost brachiopod assemblage of the San Juan Limestone (Ahtiella Zone, early Llanvirn) is characterized by a mixture of Celtic, Baltic and North American taxa, associated with few endemic and several pandemic forms (Table 1, Fig. 4). Typical genera of the Celtic Province are Ahtiella, Platystrophia, Productorthis, Inversella (Reinversella) and Rugostrophia. All of them, with the exception of Rugostrophia, also occur in the Baltic Province. On the other hand, Pomatotrema,

Taffia, Leptella (Petroria), Camerella, Idiostrophia and the recently discovered Orthidiella (unpublished data) all are distinctive Laurentian forms. In summary, of 20 genera recorded in the early Llanvirn Precordilleran limestones, 34% belong to the Celtic-Baltic provinces and 31% are Laurentian forms (ToquimaTable Head Realm). Applying correspondence

BRACHIOPODS AND PRE-ANDEAN MARGIN

65

0 "-~.............................................................

Z

50 .................................................. 40

...........................................

9 [-,

I........................................... i

!:!iii! iiiii!iiiii iiii

z

10

ARENIG

LLANVIRN

CARADOC I

Fig. 4. Variation of affinities of brachiopod faunas from the Precordillera terrane expressed in percentage of genera shared with different biogeographical provinces: AM, American Realm (Scoto-Appalachian Province); ANB, Anglo-Welsh and Baltic Provinces; B, Baltic Province; C, Cosmopolitan; Ce, Celtic Province; E, Endemics; EU, European Realm (includes Anglo-Welsh, Baltic and Bohemian Provinces); L, Laurentian taxa (Toquima Table-Head Realm); M, Mediterranean (Bohemian) Province.

analysis, Neuman & Harper (1992) concluded that the Precordilleran assemblage considered as a whole (that is, including Arenig-Llanvirn faunas), is closer to the Toquima-Table Head cluster than any of the Celtic faunas. (7) By the late Llanvirn (P. serra Zone) the only biogeographically significant evidence is from the early vertebrates, represented in the Precordillera basin by the agnathan Sacabambaspis janvieri, a species recorded in nearly coeval beds from Bolivia (Albanesi et al. 1995b). (8) Brachiopod assemblages of early Caradoc age include 17 genera (Table 2), of which only Campylorthis is restricted to the Scoto-Appalachian (Laurentian marginal) Province. The majority of taxa (seven genera, representing the 41%) belong to the European Reahn, showing affinities with both the Anglo-Welsh and Mediterranean provinces. A particularly interesting feature is the presence in the Precordillera basin of Aegiromena associated with Drabovia? and Tissintia, which indicates a strong link with the Berounian Aegiromena-Drabovia Fauna (Havli~ek 1989) of Bohemia and northern Africa (Benedetto 1995a; in press a; Benedetto & S/mchez 1996a). Of the remaining genera, two are endemic (Oepikoides and Anchoramena) and seven have a cosmopolitan or widespread distribution. Evidence from other

taxonomic groups, particularly molluscs and trilobites (Benedetto et al. 1995; Benedetto & S/mchez 1996a; Sfinchez in press; Baldis & P6the 1995), indicates that endemicity was higher during the Caradoc than during any other Ordovician epoch. Cladistic analysis of some trilobites from the early Caradoc Las Aguaditas Formation of Precordillera suggest affinities with Laurentian forms (Chatterton et al. 1997; Edgecombe et al. 1997). (9) Latest Ashgill (Hirnantian) brachiopod faunas are known from the glacigenic Don Braulio Formation (Benedetto 1986, 1990). The taxa recorded are Hirnantia sagittifera, Dalmanella testudinaria, Plectothyrella crassicosta, Cliftonia oxoplecioides, Paromalomena polonica and Eostropheodonta hirnantensis, all of which are characteristic of the Kosov Province (Rong & Harper 1988) developed in central-southern (Gondwanan) Europe, southern China, Kazakhstan, Baltica and Avalonia. Laurentian records of Kosov taxa are restricted to Perc6 and probably Maine.

Sierra de Famatina (1) Early Tremadoc (including probably latest Cambrian) times are represented by black shales

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J. L. BENEDETTO

(Volcancito Formation) containing the distinctive trilobite Parabolina (Neoparabolina)frequens and other olenids. This trilobite association is similar to those of NW Argentina, assigned by Shergold (1988) to the Baltic Province. (2) The Early Arenig shales have yielded graptolites and some trilobites including pelagic forms such as the cyclopygid Degamella (Esteban 1996). Brachiopods are very abundant in the overlying volcaniclastic sequence represented by the Suri and Molles formations (Astini & Benedetto 1996). The fossiliferous beds fall within the middle Arenig (Whitlandian, B. naris Zone) (Albanesi & Vaccari 1994; O. Lehnert, pers. comm.). The Famatinian fauna is a Celticlike assemblage with weak Bohemian affinities (Benedetto 1994). Typical genera belonging to the Celtic Province are Skenidioides, Monorthis and the recently discovered Ffynnonia, which is recorded for the first time outside Anglesey, NW Wales (Neuman & Bates 1978). Also characteristic is the genus Famatinorthis, which has recently been reported from the coeval Shin Brook Formation of Maine (Neuman 1997), a classic 'Celtic' locality. Finally it is interesting to note that specimens of 'aft. Eostrophomena sp' from the Davidsville Group of New World Island, illustrated by McKerrow & Cocks (1986, fig. 2b-f), are very similar to the Famatinian Monorthis aft. menapiae (Davidson) (cf. Benedetto 1994, pl. 1, figs 1-14). Bohemian links are suggested by the genus Prantlina, probably present both in the Suri Formation and in the Prague basin. The record of Incorthis represents a biogeographical link with the typical peri-Gondwanan faunas of Morocco (Mergl 1988), NW Argentina (Benedetto in press b) and Bolivia (Havli6ek & Branisa 1980). The record of Crossiskenidium, previously considered to be endemic to Ireland (Tourmakeady Limestone), is significant because it indicates a link with this Scoto-Appalachian locality. The occurrence of clitambonitids and pentamerids in the Famatinian volcaniclastic rocks is another feature that suggests affinity with the warm-water Precordilleran faunas. It is important to note that Famatinian brachiopods are associated with trilobites of the peri-Gondwanan Neseuretus Fauna (Vaccari et al. 1993; Vaccari 1995). These assemblages include western Gondwana (Australian) taxa, such as Hungioides and Gogoella and southern European-Avalonian genera such as Merlinia and Illaenopsis, associated with at least two endemic taxa (Famatinolithus, Pliomeridius) shared with the NW Argentina-southern Bolivia basin. The affinities between the warm-water conodont faunas of the Sierra de Famatina and Australia

were noted by Albanesi & Vaccari (1994). The peri-Gondwanan relationships are also reinforced by the record in the Famatinian basin of the bivalve Redonia, known from southern Europe, Morocco and NW Argentina (Sfinchez & Babin 1994).

Central Andean basin We include here the vast Early Palaeozoic sedimentary basin developed approximately between 10~ and 23~ in the southern Central Andes, comprising outcrops in NW Argentina, N Chile, Bolivia, Paraguay (to the west of the Asunci6n Arch) and south-central Peru (Fig. 3; Benedetto et al. 1992). (1) With the exception of a few inarticulate brachiopods, Early and mid-Cambrian faunas are unknown from this basin. Benedetto (1977) presented evidence for a latest Cambrian age for trilobite faunas from the lower part of the Santa Victoria Group of NW Argentina. Trilobites recovered from these beds, the Parabolina (Neoparabolina) frequens Zone, as well as those from the earliest Tremadoc (including Plicato-

lina, Onychopyge, Beltella, Protopeltura, Shumardia and Jujuyapis) were referred by Shergold (1988) to the cold-water Baltic Province (Olenid biofacies). This province extends across the East European, Baltic and Avalonian platforms and northern Siberia; it has also been recognized in Oaxaca (Mexico), Venezuela (El Baul Massif), Bolivia and NW Argentina. The diverse late Tremadoc trilobite assemblages of NW Argentina and Bolivia show the same biogeographical signature (Harrington & Leanza 1957). Articulate brachiopods are represented by the widespread genera Apheorthis and Nanorthis and the endemic Notorthisina (Havli6ek & Branisa 1980). (2) Arenig brachiopod faunas have been recovered mainly from the Whitlandian beds (middle Arenig) of the Cordillera Oriental and Subandean Ranges in NW Argentina and Bolivia (Havli6ek & Branisa 1980; Benedetto in press b) (Table 2). Of special interest is the occurrence in these regions of Euorthisina, known from the Mytton Flags of Shropshire, England (Williams 1974), the Moroccan Anti-Atlas (Havli6ek 1971) and Bohemia (Havli6ek 1967). Another genus in common with Morocco is Incorthis, represented in the Sepulturas Formation of Argentina by a form very close to L maroccana. The Central Andean Arenig brachiopod faunas are linked to NW African assemblages which, in turn, are related to the Bohemian faunas. Based on the associated trilobites, Waisfeld (1995) also noted provincial affinities with Avalonian

BRACHIOPODS AND PRE-ANDEAN MARGIN faunas, which is consistent with the probable presence in NW Argentina of the Anglo-Welsh genus Salopia. Brachiopod assemblages from the Sierra de Famatina and the Central Andean basin have in common four genera, of which Incorthis is of particular significance because of its more restricted geographical range (northwestern and western Gondwana). (3) Little can be said about the Llanvirn faunas, due to the almost complete absence of sedimentary rocks of that age. Tissintia has been described from Bolivia, where is associated with Aegiromena (Havli~ek & Branisa 1980), and from Peril (Hughes et al. 1980). The Bolivian species T. canalifera is very similar to T. convergens from Morocco and T. prototypa from Shropshire, and indicates close affinities with Anglo-French and Mediterranean faunas. (4) Caradoc faunas are represented in the Central Andean basin by a suite of typically Mediterranean taxa including Eorhipidomella, Destombesium, Heterorthis, Drabovinella and Hirnantia? (Havli~ek 1990; Sufirez Soruco & Benedetto 1996; Benedetto 1997a). The record of Oanduporella in supposed early Caradoc beds of Bolivia suggests a connection with Precordilleran faunas. (5) No typical members of the late Ashgill Hirnantia Fauna have been described from this basin, but some specimens assigned to Schizophoria from the Cancafiiri Formation probably belong to H. sagittifera (Benedetto et al. 1992).

67

Palaeogeographical models and biogeographical constraints A number of geodynamic models have been published in the last 15 years in order to explain the evolution of the pre-Andean margin of Gondwana and its palaeogeographical relationships with other supposedly associated continents and terranes (Herrera & Benedetto 1991; Benedetto 1993; Benedetto & Astini 1993; Astini et al. 1995; Thomas & Astini 1996; Dalla Salda et al. 1992a, b; Dalziel et al. 1994; Dalziel 1994, 1997; Keller & Dickerson 1996; Ramos 1988, 1995; Ramos et al. 1984, 1986; Bahlburg 1990; Conti et al. 1996). Each of these reconstructions fall within one of four basic geodynamic models: (a) unrelated margins, (b) paired rift margins, (c) far-travelled microplate, or (d) collision, with transfer or 'exchange' of plate fragments (Fig. 5).

Unrelated margins

In this category are the classic reconstructions in which Laurentia and the pre-Andean margin of Gondwana are placed at very different longitudes in the earliest Palaeozoic (e.g., Scotese & McKerrow 1990). Faunal migration from Laurentia to the western Gondwana margin and a similar peri-equatorial latitudinal belt have been argued to explain stratigraphical and faunal similarities during the Cambrian (Baldis &

Fig. 5. Sketch of the three palaeogeographical hypotheses for the origin of the Precordillera mountain belt (black rectangle) analysed in the text.

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Bordonaro 1982). The species-level differentiation of certain Precordilleran trilobites led Bordonaro (1989, 1992) to recognize a separate South American Province, suggesting a high degree of faunal isolation of the Precordillera at that time. Faunal evidence does not support such a reconstruction. On the contrary, there is no doubt that Laurentia and Precordillera are faunistically indistinguishable and belong to the same Cambrian palaeobiogeographical entity. In other words, by Cambrian times there is no evidence of barriers to faunal exchange separating these regions, so that they were probably geographically close.

P a i r e d rift margins

This category includes reconstructions in which western (South American) Gondwana was initially contiguous with the Laurentia plate, the posterior separation resulting from an early Palaeozoic rifting event. As a result, Precordilleran and Appalachian carbonate platforms developed as conjugate rift-margins (Bond et al. 1984; Herrera & Benedetto 1991; Benedetto 1993 model, see Fig. l b, C-D). The 'faced margins' configuration (Fig. l a; Herrera & Benedetto, 1991) represents a stage after the Neoproterozoic Pangaea configuration of Bond et al. (1984). Two presumptions characterize these scenarios: (1) the Precordillera was always a part of the Gondwana plate, which rifted away from Laurentia after the Cambrian, and (2) neither Laurentia nor the Precordillera collided with the Gondwana margin after the break-up of Pangea. Biogeographical evidence supports a direct connection only between Laurentia and the Precordilleran region of South America, not with the entire 'Pacific' margin of Gondwana such as depicted by Bond et al. (1984). Much uncertainty exists about the affinities of the early-mid-Cambrian faunas developed outside the Precordillera. The only biogeographically significant locality is in the Sierra de La Macarena, Eastern Cordillera of Colombia, where a trilobite association including the genera Asaphiscus, Ehmania and Paradoxides was recorded (Harrington & Kay 1951; Rushton 1963). The last genus is distinctive of the Acado-Baltic Province and its record in North America is restricted to the Carolina and Avalon accreted terranes (Secor et al. 1983, N e u m a n et al. 1989). The complete absence of Paradoxides in the autochthonous basins of North America makes unlikely the juxtaposition of Laurentia and the northern Andean margin of Gondwana in the mid-Cambrian.

This assumption is supported by the 'European' (Baltic Province) biogeographical signature of the latest Cambrian-Tremadoc trilobite faunas founded from NW Argentina to the E1 Bafil Massif of Venezuela, even if the palaeogeographical significance of such olenid faunas has been seriously questioned (Fortey & Cocks 1986). The pattern of Cambrian to earliest Ordovician trilobite affinities indicate that a sharp faunal boundary exists between the Precordillera and the remaining Andean basins. The record of Asaphiscus in the Lizoite Formation of the Central Andean basin mentioned by Acefiolaza & Bordonaro (1988) needs confirmation by further studies because the poor preservation of this material prevents an accurate generic identification (Vaccari pers. comm. 1996). Such a biogeographical boundary also coincides with strong contrasts in the lithological succession, the Precordillera being characterized by carbonate rocks, whereas the Famatina and Central Andean basins are entirely filled by siliciclastic rocks.

Far-travelled microplate

According to this model (Fig. lb, A-B), the Precordillera and surrounding 'Grenvillian' metamorphic rocks represent a far-travelled microplate (Precordillera, or Cuyania, terrane) rifted off from probably the southeastern margin of Laurentia and later accreted to the Gondwana margin (Benedetto 1993; Benedetto & Astini 1993; Benedetto et al. 1995; Astini et al. 1995, 1996; Thomas & Astini 1996; Ramos 1995). This reconstruction is consistent, not only with the identity of Cambrian trilobites from both regions, but with the progressive change in faunal provinciality of the Precordilleran brachiopods through the Ordovician (Fig. 4). Although the rifting phase probably started in the Early Cambrian, the Precordillera became faunistically distinguishable from Laurentia only from the earliest Arenig. At that time, and especially during the mid-late Arenig and early Llanvirn, several brachiopod genera unknown from the Laurentian carbonate platform appear. By a trans-latitudinal trajectory across the southern Iapetus, the microcontinental model provides an elegant explanation for the faunal exchange between the Cuyania terrane and the northern Iapetus volcanic islands and subduction-related volcanic arcs inhabited by Celtic faunas (Neuman & Harper 1992). The strong Baltic signature of the Precordilleran brachiopods remains ambiguous because in our reconstruction (Astini et al. 1995, fig. 14c) the Baltica

BRACHIOPODS AND PRE-ANDEAN MARGIN palaeocontinent is too far away from the CuyaDia terrane. This is also the case of the northern Iapetus volcanic islands inhabited by 'Celtic' brachiopods (Neuman 1984). Estimation of the distance at which the Precordillera was able to attain faunal distinctiveness largely depends on the inferred effectiveness of larval migration and patterns of palaeoceanic circulation within Iapetus. Assuming a current velocity of c. l k m / h - : and a planktonic larval stage of 6 weeks, the maximum distance at which 'Laurentian' larvae

69

could reach the Precordillera, or vice-versa, is c. 1000-1500km. This range is consistent with separation at a standard sea floor spreading rate during c. 50 Ma (i.e., Early Cambrian to Early Arenig). An unanswered question is why Laurentian forms such as Syntrophia, Nothorthis, Cuparius, Leptella (Petroria) and Oligorthis migrated from Laurentia to the Precordillera, but none of the Celtic, Baltic and endemic genera of the Precordillera colonized the Appalachian carbonate platform. It seems that this, as well as the strong faunal exchange between the

L

oo,, n

12 ,9 .... ,

+

[

uric

i

Early Ordovician --] (Late Arenig-Eady L!anvirn)l

9

-.

1000 km

Fig. 6. Early Ordovician schematic palaeogeographical reconstruction of the Southern Iapetus. C, Inferred original location of the Carolina terrane; EA, Eastern Avalonia; EP, Eastern Sierras Pampeanas ; F, Famatina volcanic arc; FEP, Puna volcanic belt (includes the 'Faja Eruptiva de la Puna Occidental' and 'Oriental'); GB, Guyana/Brazilian craton; KN, Kola nappe; P, Precordillera terrane (Cuyania); Punc, Puncoviscana folded belt; T, Northwestern Ireland terrane (includes Tourmakeady Limestone.); WA, Western Avalonia.

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Precordillera terrane and the northern Iapetus islands, can only be explained by ocean currents controlling the larval dispersal (Benedetto et al. in press). The timing of Precordillera accretion to Gondwana is still poorly constrained. Originally we assumed that by the Arenig, a narrow seaway separated the Precordillera terrane from the Famatinian arc-related volcanic rocks, and that docking of the Precordillera with Gondwana occurred during the Llanvirn (Benedetto & Astini 1993). The Famatinian faunas combine taxa from the Gondwana continental shelf and Celtic and Laurentian forms from the Precordillera terrane. This two-way colonization is what might be expected if the Famatinian volcanic arc lay, as we suppose, between Gondwana and the Precordillera (Fig. 6). The first 'Gondwanan immigrant' into the Precordillera is the shallow-water agnathan Sacabambaspis, recorded in the Llandeilo, followed in the Caradoc by brachiopods of the Mediterranean (and Gondwanan) DraboviaAegiromena Fauna. At this time, Cuyania received only a few Laurentian colonizers and attained the highest degree of endemism since its separation from Laurentia. The relatively high proportion of strophomenides and the AngloBaltic signature of several taxa, together with the persistence of carbonate deposits, suggest warmer waters than those in the autochthonous Gondwanan basins (Benedetto & Sfinchez 1996b). By Caradoc times, the Precordillera was probably not yet accreted or, if it was, some kind of barrier separated it from the Central Andean basin. Palaeontological evidence for the accretion is conclusive only from the early Wenlock, when the typical Clarkeia Fauna (Afro-South American Realm) became widely distributed across the basins of southern South America (Benedetto & Sfinchez 1996b). The widespread late Ordovician Hirnantia Fauna, well documented in the Precordillera, does not necessarily indicate geographical continuity with Gondwana; but merely, because of its association with glacigenic horizons, suggests a mid- to highlatitude position at the end of the Ordovician.

CoO&ion with exchange o f plate fragments The reconstructions included here differ significantly from the preceding ones, in that the Precordillera was transferred to the pre-Andean margin through Laurentia-Gondwana collision followed by new rifting. Such a rift system was not exactly coincident with the former collisional suture, so that the Precordillera remained

attached to Gondwana in a similar way to that by which the eastern Newfoundland was transferred from Baltica to Laurentia when Iapetus closed and re-opened (the Wilson Cycle archetype). Two different models have been proposed, which basically differ in the size and location of the transferred 'allochthonous' area in relation to the Laurentia plate: (1) the 'Occidentalia' model and (2) the 'Texas Plateau' model. The first hypothesis, based in the Cambrian palaeogeographical reconstruction of Dalziel (1991), was proposed by Dalla Salda et al. (1992a, b), who interpreted the Famatinian orogen as the result of the Laurentia-Gondwana collision, with the Occidentalia terrane representing a continental sliver detached from eastern North America following the collision. According to this reconstruction, the Precordilleran carbonate belt represents the continuation in South America of the Appalachian platform. If correct, faunal exchange between the two areas must have been free and continuous at least until the late Ordovician rifting event. However, it is now clear that biogeographical divergence increased after the early Ordovician, and that by the mid-Ordovician the Precordilleran faunas became markedly different from those of the Appalachians. This and the latitudinal incompatibility of climatically-sensitive sediments (late Ordovician carbonates and evaporites in Laurentia versus glacigenic diamictites in Gondwana) leads to rejection of this reconstruction (Benedetto 1993). Recently, Dalziel (1997, p. 34) pointed out that the 'Occidentalia' model of Dalla Salda et al. (1992a, b) contradicts the faunal and palaeoenvironmental differences between Gondwana and Laurentia during the Ordovician. According to his 'Texas Plateau' hypothesis (Dalziel 1997), the Precordillera was part of a continental plateau outboard of the Ouachita embayment, whose tectonic, geodynamic and sedimentary history was analogous to the present-day 'Malvinas/Falkland Plateau'. This Laurentian promontory was originally rifted from Gondwana near the Proterozoic-Cambrian boundary, subsequently collided with the proto-Andean margin by the mid-Ordovician forming the 'Artejia' megacontinent (Laurentia plus Gondwana) and finally detached from Laurentia during a Caradoc rifting event. Although this hypothesis attempts a better integration of plate motions and faunal data than the 'Occidentalia' model, a number of biogeographical objections can be made to this scenario: (1) If, as Dalziel (1997, fig. 15) postulates, the Precordillera was a part of the 'Texas continental plateau' by the early Ordovician (475 Ma), then a nearly complete identity might

BRACHIOPODS AND PRE-ANDEAN MARGIN be expected between the 'Laurentian Plateau' faunas and those on the Laurentia plate (Great Basin, Appalachians). This is based on the fact that the benthic faunas of both the Malvinas/ Falkland plateau (the Mesozoic-Tertiary analogue) and the contiguous Patagonian shelf belong unequivocally to the Magellanic Province (Ringuelet 1954; see also Valentine 1973). However, in Early Ordovician (Canadian and Early Whiterock) times, Precordilleran brachiopods are markedly different from those of Laurentia: as stated above, only 31% are typical ToquimaTable Head genera while the other associated forms are from other realms or endemic. (2) After the mid-Ordovician collision and the subsequent closure of the 'Western Iapetus', a bridge of about 1000kin of submerged continental crust was established between Laurentia and Gondwana (Dalziel 1997, fig. 16). If correct, the progressive interchange of brachiopods along this migratory route should have produced a break-down of provinciality. In contrast, there are in reality few mid-Ordovician genera in common between the Precordillera and Laurentia and between the Precordillera and Gondwana. Likewise, in spite of the supposed relative proximity of both continents by the early Caradoc, no Gondwanan brachiopods have ever been recovered from Laurentia. In summary, from a dynamic perspective, the 'Texas Plateau' model only explains the increasing affinities between Precordilleran and Gondwanan faunas through the Ordovician, but does not account for the corresponding increase in the faunal differences between the Precordillera and Laurentia.

The Precordillera, a microplate or a Gondwanan plateau? From the faunal evidence summarized above it is clear that in Cambrian times, the Precordillera and Laurentia were geographically close. With the exception of the 'unrelated margins' hypothesis, all of the proposed reconstructions agree on this point. The fact that Laurentian and Precordilleran faunas diverge after the early Ordovician leads to the conclusion that the Precordillera gradually moved away from Laurentia as an independent continental plate and oceanic-crust floored sea-way separated both plates. It should be noted, however, that such faunal divergence is also biogeographically compatible with the 'paired rift margins' hypothesis. In the Benedetto (1993) scenario (reproduced in Fig. 1b, C), the Precordillera is represented as a Gondwanan promontory which could have

71

been genetically similar to the Laurentian 'Texas plateau' envisaged by Dalziel (1997). From the data reviewed here, the only palaeontological support for this reconstruction is the relatively high percentage of shared forms with the Precordillera: of the 13 genera identified in the Famatina basin, 7 (54%) have also been recorded from coeval beds in the Precordillera, suggesting a closer relationship between both regions than was previously accepted. Moreover, the Famatina brachiopod fauna seems intermediate between the mixed LaurentianCeltic associations of the Precordillera and the nearly pure Gondwanan ('Mediterranean') faunas of the Central Andean basin. It is clear that a faunal gradient existed from the Central Andean to the Precordillera basins. The problem is to determine if such a gradient was the result of a gradual climatic change from temperate to warm waters along a nearly continuous continental margin across palaeo-latitudes (a geographical and climatic cline) or, in contrast, caused by other factors such as oceanic barriers and/or a greater geographical separation than that seen today. Because of its analogy with the Ordovician basins, a better example of climatically controlled facies and faunas is that of the autochthonous Permian basins of Gondwana. Warm-water carbonate facies are widely distributed from Venezuela to southern Bolivia (Titicaca Group), approximately between the Permian palaeo-equator and 30~ These Central Andean successions have yielded highly diverse, warm water brachiopod faunas belonging to the Texas Province of the Tethyan Realm (Shi & Waterhouse 1991). Conversely, coeval rocks of the Precordillera basin (Paganzo Group) are nearly exclusively clastic and contain temperate to cold-water brachiopods of the Gondwanan Realm. Moreover, the Tepuel basin (Argentine Patagonia) and the ParaM basin (Brazil), located to the south and east respectively of the Precordillera basin, both contain early Permian glacigenic sediments. Present-day separation between the Central Andes and the Precordillera is about 12~ of latitude, and between the former and the Patagonian and Brazilian basins reaches about 18-20 ~. According to this analogy, the early Ordovician faunal-lithological gradient could be explained by means of a latitudinal separation like that inferred in the early Permian model, which is comparable to the present-day. To account for the absence of warm water indicators in the early Ordovician rocks of NW Argentina and Bolivia, the Precordillera would have to have been located at about 30 ~ and the

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Central Andean basin at c. 40-45 ~ Waisfeld & Astini (1993) pointed out that both sedimentary regimes, siliciclastic in NW Argentina and carbonate in the Precordillera, could have coexisted along the same Gondwanan margin, as occurs in the Recent southern coast of North America, where the Florida peninsula separates deltaic sedimentation to the west and carbonate deposition to the east. Although the 'Gondwana Plateau' model seems to be consistent with development of the three regions (Central Andes, Famatina, Precordillera) along the same pre-Andean continental margin, the following evidence favours the 'microplate' model. (1) No Cambrian trilobites of Laurentian signature have been encountered in the autochthonous pre-Andean basins. Certainly not enough is known yet about Early-Mid Cambrian faunas, but those recorded from Colombia, as stated above, indicates European affinities. (2) Latest Cambrian and Early Tremadoc trilobite associations from the outer shelf black shales of NW Argentina and coeval units in Bolivia are represented by 'Olenid biofacies' (Fortey 1975) which are generally thought to be independent of palaeogeography. However, the shallow water facies have also yielded only Avalonian and Baltic trilobites. Likewise, shallow-water communities from late Tremadoc beds of the Central Andean basin also lack Laurentian trilobites. (3) The high endemism exhibited by the Arenig brachiopods of the Famatina basin, as well as petrological evidence, is compatible with a volcanic arc system rather than an epicontinental basin. (4) Differences in tectono-stratigraphical history preclude the geographical continuity of the Famatina and the Precordillera during the Cambrian-mid Ordovician time interval and suggest that they were somewhat isolated by oceanic barriers and probably more separated than at present. In summary, although the 'Gondwana plateau' is consistent with the distribution of climatically sensitive sediments and, at least partially, with the early Ordovician faunal evidence, the 'far-travelled microplate' model fits better with all the available palaeontological and geological data.

The Ordovician Famatina-Puna-Avalonia volcanic island-arc system Since the seminal synthesis of northern Iapetus geology and biogeography by Neuman (1984), Ordovician volcaniclastic successions within the

Appalachian-Caledonide orogen are generally interpreted as insular in origin. Such volcanic islands are thought to have developed at midto high latitudes, in both mid-oceanic and peri-continental settings. From a biological viewpoint, they were inhabited by distinctive brachiopod assemblages, constituting the Celtic Province characterized by a relatively high proportion of endemics (c. 46% according to Neuman 1984), some genera that occur later on the Laurentian platform, and several Baltic forms. These Celtic assemblages were first described from Arenigearly Llanvirn rocks of Anglesey and the Irish Rosslare terrane and further recognized in volcaniclastic rocks from Central Newfoundland, Maine, New Brunswick, Norwegian Caledonides, the Argentine Precordillera, and probably NE China (Neuman & Harper 1992 and references therein). The new palaeontological evidence indicates that the Famatina volcanic sequence must also be included in the Celtic Province. According to the interpretation of Neuman and Harper (1992, fig. 3) this developed on the periphery of Gondwana, whereas the Toquima-Table Head Realm (North American) represented peri-equatorial passive-margin settings and related low-latitude terrains. Petrological and geochemical evidence suggest that the Famatina System was a magmatic arc related to an active margin (Acefiolaza & Toselli 1988; Toselli et al. 1993). These volcanic rocks were interpreted by Manheim (1993) as islandarc complexes and the associated volcanosedimentary succession as a back-arc setting. Genesis of this volcanic chain seems related to the gradual closure of the ocean separating the Precordillera and Gondwana. We previously speculated that east-dipping subduction was initiated in the Tremadoc and ended near the Arenig-Llanvirn transition with the accretion of the Precordillera terrane (Benedetto 1993, Benedetto & Astini 1993). The back-arc basin was filled by a c. 2000m thick succession (Suri and Molles formations) in which Astini & Benedetto (1996) recognized a first stage of rapid thermal subsidence, characterized by starved-basin graptolitic black shales, followed by a shallowing upward platform succession punctuated by storm layers and prograding volcaniclastic wedges (Mfingano & Buatois 1994). Deposition culminated with shallow-water reddish siltstones and sandstones interbedded with volcanic breccias and fossiliferous tufts. The mid-Arenig volcaniclastic rocks of the Famatina basin are richly fossiliferous but unfortunately their study is not concluded. However, their brachiopod faunas are characterized by a dominance of Celtic forms associated

BRACHIOPODS AND PRE-ANDEAN MARGIN with few typical Gondwanan genera and at least four endemics. Particularly relevant to the biogeographical interpretation is the presence of the Welsh genus Ffynnonia in the Suri Formation and of Famatinorthis in the Shin Brook Formation of Maine (Neuman 1997). This new palaeontological information confirms our inferred palaeogeographical connection between the Famatina island-arc complex and northern Iapetus islands and microplates bearing Celtic brachiopods (cf. Astini et al. 1995, fig. 14b). The Famatina System probably continues farther north into the Puna Region of NW Argentina, Chile and Bolivia, where thick volcano-sedimentary successions consisting of rhyolitic lavas, siliceous tuff breccias, spilites and volcaniclastic turbidites are recorded ('Faja Eruptiva de la Puna' volcanic belt, Fig. 3). Associated fossils are Tremadoc to Arenig in age; Arenig beds have yielded the trilobites Annamitella, Platillaenus and Illaenus (Koukharsky et al. 1996) and several articulate brachiopods which have not been studied. This volcanism has been related by several authors (Coira et al. 1982, Koukharsky et al. 1988) to an east-dipping subduction zone which was active until the late Arenig. Recent palaeomagnetic measurements from the early Ordovician rocks of Famatina and Puna Oriental (Conti et al. 1996) provide independent evidence that these two regions may constitute a separate terrane located near the Gondwana margin and geographically related to the early Ordovician Central Newfoundland volcanic sequence. The peri-Amazonian placement of West Avalonia is also supported by Nd-isotope data and the age of detrital zircons (Nance & Murphy 1994). To account for this extended F a m a t i n ~ P u n a Avalonia volcanic arc, a nearly continuous subduction zone is required along the preAndean Gondwana margin, similar to the Puna-Arequipa-Perij~ east-dipping subduction zone inferred by Scotese & McKerrow (1991). Another continental fragment involved in the Iapetus Ocean is the Irish Northwestern terrane. Shelly faunas from the Arenig Tourmakeady Limestone described by Williams & Curry (1985) as well as those from late Llanvirn beds (Harper & Parkes 1989) can be assigned unequivocally to the Toquima-Table Head Realm. The Tourmakeady Limestone assemblage is remarkable because of 23 articulate brachiopod genera, 9 (39%) are endemics. This unusually high endemism strongly suggests a peri-insular setting for the Tourmakeady fauna. One of these Irish 'endemics' (Acanthotoechia), has been recently discovered in Arenig limestones near Talacasto,

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central Precordillera, and another (Crossiskenidium) seems to be represented in the Suri Formation (mid-Arenig) of the Sierra de Famatina. The Tourmakeady and San Juan formations have 12 genera in common, which represent 52% of the Irish fauna. On the basis of these affinities, it might be inferred that the Northwestern terrane of Ireland, which is usually placed near the northernmost Laurentian margin (cf. Neuman & Harper 1992), could have been geographically close to both the Precordillera terrane and the Famatina island arc. Inferred positions of the Precordillera, Famatina, Puna, east and west Avalonia and related terranes are shown in Fig. 6.

Iapetus geodynamics and the early evolution of the pre-Andean margin L a t e Proterozoic to Early Ordovician

Assembly of palaeocontinents, microplates and volcanic islands prior to the early Palaeozoic opening of the ]apetus Ocean is a difficult task because of the absence of palaeontological criteria. On the basis of the biogeographically better constrained early Palaeozoic reconstructions, however, some inferences about Neoproterozoic palaeogeography can be made by means of back-stepping analysis. The first point concerns the identification of the Laurentian margin counterparts. Since the hypothetical reconstruction of Bond et al. (1984) and the assembly proposed by Hoffman (1988), the preAndean margin of South America seems a better candidate than that of northwestern Africa. In the late Proterozoic (750Ma) reconstructions made by Hoffman (1991) and Dalziel (1991), eastern Laurentia is juxtaposed with Amazonia, and Ramos et al. (1993) add the Pampia and Patagonia plates. All of these continental blocks are thought to have been amalgamated during the Grenville orogeny (c. 1.1 Ga), even if the final assembly of Gondwana did not occur until the Brazilian (Pan-African) orogeny. Dalziel (1993, 1994) suggested, based on geochronological evidence, that the conjugate margin of the Hebridean shield of NW Scotland and Ireland may have been the South American Arica embayment. Both the Arequipa-Antofalla massifs and the Labrador-Scotland-Greenland promontory have yielded 'old' isotopic ages of c. 1.8-1.9 Ga and 'Grenvillian' ages. On the basis of the similarities between the brachiopod faunas of northwesternmost Ireland (South Mayo Trough) and the Precordillera, I suggest that both terranes could have been relatively close during the early Ordovician and that this

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geographical proximity may have been due to a similar history of rifting and drifting from the Laurentian margin. In the reconstructions above, Laurentia rifted from Gondwana by the end of the Proterozoic. Development of the Appalachian passive margin is well documented by a complete succession of late Precambrian rift-facies, followed by clastic and carbonate sediments recording transition to the drifting phase in the early Cambrian (Hatcher 1989 and references therein). On the pre-Andean margin, no conclusive evidence of a coeval rifted margin has been encountered. The thick turbidite-like sequence of the Puncoviscana Formation and its associated bimodal magmatism is interpreted as being developed on a passive margin facing the Iapetus (Ramos 1988) or as an intracratonic basin developed on thinned crust (Omarini & Sureda 1993). The Puncoviscana rocks contain a varied ichnofauna of Vendian, Tommotian and earliest Cambrian age, the latter dominated by Oldhamia (Acefiolaza & Durand 1986). Late Precambrian to early Cambrian Oldhamia-bearing slates and sandstones (Grand Pitch Formation) crop out in the Gander terrane of the northern Appalachians. These rocks were deformed by the Late Cambrian Penobscot orogeny, which has also been recognized along the Scottish-Irish Caledonides (Grampian orogeny) and in the Scandinavian Caledonides (Finmarkian orogeny) (Neuman &

Max 1989). It is important to remember that metamorphic rocks of the Gander terrane in the Penobscot County of Maine are unconformably overlain by the volcano-sedimentary Shin Brook Formation in which Famatinorthis occurs (Neuman 1997). Another allochthonous terrane to the east of the former Appalachian margin is the Carolina terrane, composed of thick marine metasedimentary and volcanic rocks, probably deposited on a thin continental crust (Horton et al. 1989). Fossils include soft-bodied Ediacaran-like metazoans and the trilobite Paradoxides (Secor et al. 1983). The Carolina slate belt rocks were affected by Late Precambrian-Early Cambrian? deformation, but the age of regional metamorphism suggests that the Carolina terrane was accreted to Laurentia during the Taconian orogeny (Horton et al. 1989). The Avalon terrane of New England and eastern Newfoundland also has yielded early-mid-Cambrian trilobites similar to those of the Carolina terrane. Palaeontological evidence suggest that these Appalachian terranes have stronger affinities with peri-Gondwanan terranes than with Baltica (Neuman et al. 1989). Although the Neoproterozoic geology is highly complex and the available evidence is not enough to propose an accurate reconstruction, the data reviewed above suggest the following (Fig. 7).

"'--.

anl

1000 km -

Fig. 7. Schematic palaeogeographical reconstructions for the early opening of the Southern Iapetus Ocean. Late Precambrian map, modified from Dalziel (1994).

BRACHIOPODS AND PRE-ANDEAN MARGIN (1) The low-grade Vendian-Early Cambrian Puncoviscana Formation passive margin deposits (Je~ek & Miller 1987) and the broadly coeval similar rocks of the allochthonous Carolina and Gander terranes were probably deposited in a narrow, rapidly subsiding basin developed between Laurentia and Gondwana during an early rifting phase (cf. Dalziel et al. 1994) (2) After break-up, Laurentia moved from high to low latitudes. The Precordillera carbonate platform and the Northwestern Irish terrane developed on the eastern margin of Laurentian. (3) Both microplates detached from Laurentia, probably at different times: initial rifting of the Precordillera began in the early Cambrian, but final separation from the Ouachita embayment did not occur until the mid-late Cambrian (4) Folding and metamorphism of the Puncoviscana Formation and associated plutonism (Tilacaric orogeny), which is by far the most important tectonic event in the Central Andean basin, as well as the broadly coeval Penobscot orogeny, could be related to the start of the eastward subduction of the Iapetus oceanic crust beneath the Gondwana plate. The absence of such a disconformity in the Precordillera basin is another indication that they were not on the same margin in the Early-Mid-Cambrian. As Miller (1996) has pointed out, the change

/

/

75

from a passive to an active margin occurred when Laurentia was far away from the Gondwana margin (5) This subduction zone probably extended along the greater part of the pre-Andean margin and continued northwards into northwestern Africa and the margins of Baltica. When subduction began, arc-related volcanism developed extensively along the active margin, forming the Famatina-Puna-Avalonia volcanic arc. It is conceivable that the Late Arenig to early Llanvirn K-bentonites recorded in the Precordillera basin resulted from this volcanic activity (Huff et al. 1995) (Figs 6, 7).

Late Ordovician

In the Southern Andes, the Precordillera gradually approached the Gondwana margin before final accretion. However, faunal evidence cannot provide constraints on the timing of the collision. It seems likely, however, that by the Caradoc the Precordillera was near, or accreted to, the Famatina arc, resulting in closure of the back-arc basin (Fig. 8). This is also reflected in a rapid change in sedimentary facies of the Precordillera, with drowning of the carbonate platform and widespread black shale sedimentation (Astini et al. 1995). In the Puna region,

S

/ //

JVA

/

///////////"

//

/ i/

//

// // /

oooo \\

/

// /

k

/

t

/

/

/

/

./

/

Late~

"'~'0o

LateSilurian 1 /

Fig. 8. Inferred Laurentia and Gondwana interactions during the Late Ordovician and Silurian. Key as in Fig. 6.

76

J. L. BENEDETTO

volcanic activity ceased when the ArequipaAntofalla terrane collided with the Gondwana margin transforming the Puna back-arc into a foreland basin (Palma 1990; Bahlburg 1990: Bahlburg & Breitkreutz 1991). During the Arenig-Llanvirn, Laurentia moved along the equatorial belt around Gondwana from the 'western' to the 'eastern' Iapetus. The main modification that I introduce to the Dalziel (1991) wander path is that the site of the Late Ordovician collision between the eastern margin of Laurentia and the Avalonian island arcs is the northwestern corner of South America (Fig. 8) instead of the opposite margin of the 'Pacific' Ocean (cf. Dalziel 1997, fig. 17). This is supported by the following: (a) The Taconian orogeny along the Appalachian fold belt has been related to accretion of terranes (e.g., Hatcher 1987). One, the Carolina terrane, is generally thought to have originated in a volcanic-arc setting above an east-dipping subduction zone (Hatcher 1989); the age of metamorphism suggests that it was accreted during the Taconian orogeny (Horton et al. 1989). (b) Because of the strong lithological similarity to the Puncoviscana belt and their Atlantic fauna, it is not unreasonable to locate the Carolina terrane as bordering the Gondwana margin, as was suggested by Dalziel (1997), and near the localities with the Paradoxides fauna (La Macarena Massif and Cordillera Oriental of Colombia) (Bordonaro 1992). The Late Ordovician palaeolatitude of about 22 ~ suggested for the Carolina terrane by Vick et al. (1987) is consistent with the presence of mid-Cambrian carbonate rocks in Colombia. Therefore it is conceivable that transfer of this terrane to Laurentia occurred as the result of the late Ordovician collision. (c) Similarly, the late Precambrian volcanosedimentary succession and overlying stableplatform shales and sandstones with an Atlantic Cambrian fauna in the Avalon Zone of Newfoundland and related zones of New Brunswick might have formed along the NW Gondwana margin and been transferred by the same mechanism. (d) There is strong evidence to support a 'Caledonian' orogeny in NW South America, including widespread magmatism of 475 Ma to 439 Ma in the Cordillera Oriental and the Santa Marta Massif of Colombia as well as in the M6rida Andes of Venezuela (460 Ma to 430 Ma) (Gonz/tlez de Juana et al. 1980 and references therein). Stratigraphical units of the Sierra de Perijfi and Cordillera Oriental of Colombia include early Palaeozoic, low to high-grade

metamorphic rocks unconformably overlain by undeformed Early Devonian sedimentary rocks (Benedetto 1992; Pimentel de Bellizzia 1992). I also infer that the volcaniclastic Shin Brook Formation of Maine and the volcanic rocks of central Newfoundland developed originally on the periphery of NW South America and were transferred to Laurentia during the Taconian collision and the subsequent break-up. If the NW African margin is replaced by the NW corner of South America, the reconstruction proposed here is very similar to that presented by Pickering et al. (1988), based on a review of palaeomagnetic, petrological, stratigraphical and structural evidence. According to this model, by the Late Ordovician, Eastern Avalonia and Laurentia were geographically close and the Iapetus Ocean was nearly closed, but with a narrow sea-way persisting until the end of the Silurian. Faunal exchange between the Mediterranean and the Central Andean basins during the Caradoc might have occurred along the epicontinental sea developed outboard of the West Africa, Guayana and Amazon cratons.

Silurian and Devonian

The distribution of post-Taconian sedimentary rocks in the northern Appalachians is unrelated to the early Palaeozoic tectono-stratigraphical zonation. Silurian and Devonian rocks have been interpreted as deposits of successor basins formed across the vestiges of the Iapetus margins (Williams 1979). Authors such as Arnott et al. (1985) and Pickering et al. (1988) have suggested that late Caradoc to Wenlock basins were bounded and controlled by thrusts in the Dunnage zone, whereas in the Gander zone lithospheric loading led to the development of a deep marine foreland system. Evidence from brachiopod distribution can be used to test the scenario proposed here (Fig. 8). Brachiopod faunas from the E1 Horno Formation of the M&ida Andes, northern Venezuela, were described in detail by Boucot et al. (1972). These faunas can be unequivocally assigned to the North Atlantic Region of the North Silurian Realm. In particular, the late Llandovery-early Wenlock Antirhynchonella Fauna shows close relationships with the Back Bay fauna of southern New Brunswick (Boucot et al. 1966). Recently Boucot et al. (in press) stated that the Venezuelan Antirhynchonella Community has a slightly more European than North American aspect. I suggested that during the Silurian, the Venezuelan basin was closely related to the

BRACHIOPODS AND PRE-ANDEAN MARGIN Appalachian-Avalonian basins (Benedetto 1984, fig. 39). On the other hand, the record of Eocoelia sp. in the Llandovery of Paraguay (unpublished data) and the discovery in the Precordillera of several taxa showing strong affinities with the earliest Llandovery faunas of Wales (Benedetto 1997b) suggest that, as in the Late Ordovician, a faunal exchange occurred along the marine platforms surrounding the Guyana and Amazon cratons. A gradual isolation of Bolivian, Argentinian and especially of the intracratonic Brazilian basins, however, led to the differentiation of the cold water, AfroSouth American brachiopod Realm (Benedetto & Simchez 1996b). Provincialism of the highly diverse Devonian brachiopod faunas from the Sierra de Perij'fi of Venezuela provides better evidence to support the juxtaposition of Laurentia and NW South America (Benedetto 1984). It is proposed that such a continental assembly during the Devonian was inherited from the Late Ordovician/ Silurian ('Caledonian') continental collision previously suggested. Of the ten species recorded in the Carlo Grande Formation (Siegenian) of Venezuela, nine are identical to species from the Oriskany Formation of the Appalachian basin. Likewise, the Emsian-Eifelian-Givetian rocks of the Sierra de Perijfi (Rio Cachiri Group) have yielded more than 50 genera, all of which are also present in the Appohimchi Subprovince of the Eastern Americas Realm. Moreover, herbaceous lycophyte plants from the upper part of the Venezuelan succession show the greatest similarity to those from New York State (Edwards & Benedetto 1985; Berry et al. 1993). Similarities with the Appalachians are not only evident in the palaeontological content but also in the stratigraphical sequence: the succession of Carlo Grande (shallow water sandstones and biostromal limestones), Carlo del Oeste (platform shales and siltstones) and Campo Chico (prograding delta front sandstones) formations is nearly identical to the Oriskany-Onondaga-Marcellus-Catskill sequence of New York State, including the record of volcaniclastic rocks coeval with the Tioga Bentonite. To account for these facts, I postulate that the sedimentary basins of the Central Appalachians and NW Colombia-Venezuela were geographically continuous. The resulting reconstruction (Benedetto 1984), reproduced here in Fig. lc, is very similar to the palaeogeographical model by Dalziel et al. (1994) using palaeomagnetic data from Van der Voo (1993, fig. 4). In relation to the Devonian assembly, Dalziel et al. (1994) pointed out that the Acadian orogeny involved not only arc collision but also strike-slip movements. Similar

77

displacements were also detected in Central Newfoundland (Pickering et al. 1988). Assuming that such dextral movement continued during the Carboniferous, placement of the NW corner of South America at or near the southern margin of North America (approximately in the present-day Gulf of Mexico) can be expected in the late Palaeozoic, as is depicted in current palaeomagnetically controlled reconstructions. A corollary of the plate interactions proposed in this paper is that from Late OrdovicianSilurian times, Gondwana, Laurentia, Baltica and Avalonia became merged in a unique continent. To the south of the present-day Gulf of Guayaquil, the pre-Andean margin opened onto a wide ocean at least partially analogous to the Pacific Ocean, for which I propose the name 'Notus Ocean' (Fig. 8).

Concluding remarks The high level of provincialism displayed by articulate brachiopods through the major part of the Palaeozoic Era provides invaluable evidence to establish spatial relationships between macroand micro-plates and is therefore a useful tool to test different palaeogeographical hypotheses. However, the criterion of concordance of biological evidence with independent geological and geophysical data is the only way to construct satisfactory palaeogeographical maps. The biogeographical affinities of Early Palaeozoic brachiopods and associated shelly faunas published in the last 15 years from western Argentina, Bolivia, and Venezuela suggest that the preAndean margin of Gondwana was the site of complex plate interactions. One of the most important conclusions is that the history of this margin was linked to the evolution of the Iapetus Ocean, formerly restricted to the present-day North Atlantic continents. It is important to remark that there is significant agreement between the biogeographically-controlled maps proposed here and the sequential set of reconstructions proposed by Dalziel in successive papers since 1991, but in particular with his last interpretation of 1997. By different methods, and based on qualitatively different data, both of us have suggested that East Laurentia faced western South America, especially during the Cambrian and Ordovician, and that some kind of interaction between the two continents was established during the Devonian Recently, Dalziel (1997) stated that the 'microcontinental' model of Benedetto (1993) does not account for the relative positions

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of Laurentia and Gondwana, concluding that 'In the case of Precordillera, if attached to Laurentia at the time of Ordovician docking in Gondwana, it is a palaeogeographical key, defining the relative palaeolongitudes of those major continental entities at certain times prior to the amalgamation of P a n g e a . . . ' . I agree completely with this possibility, but the critical question is what evidence is there to support the idea that the Precordillera was in such a position at the time of the collision? Trilobite evidence, as it is known, indicates that the Precordillera carbonate platform developed on the eastern margin of Laurentia. The subsequent Ordovician history of the Precordillera terrane can be now reconstructed accurately on the basis of a very substantial amount of palaeontological data, including not only articulate brachiopods but sponges, bryozoans, bivalves, ostracods, trilobites and graptolites. In this respect, evidence seems to be conclusive that the Precordillera moved away from Laurentia as an independent plate. Change of faunal affinities through time, from exclusively Laurentian in the Cambrian to dominantly Gondwanan in the Silurian, with intermediate 'mixed' LaurentianCeltic-Gondwanan associations during the Ordovician does not support, in my opinion, the hypothesis that the Precordillera was a Laurentian plateau until the mid-Ordovician. In general, displacement of the Precordillera fits well in the theoretical model (case C) proposed by Fortey & Cocks (1986) in which two opposing continental margins lie at different latitudes and the 'island' occupies an intermediate latitudinal and geographical position. Until recently, it has generally been assumed that the East Laurentian terranes accreted during the TaconianAcadian orogenies, such as Avalonia, Indian Bay, Maine, and the volcanic sequences of Central Newfoundland, were originally located near the northern edge of Gondwana, relatively close to Baltica, forming microcontinents, volcanic arcs or intra-Iapetus islands (e.g., the Central Newfoundland volcanic rocks) (Neuman 1988). However, new palaeontological evidence suggests that, by the Arenig, 'North Iapetus' and the Famatina volcanic sequence bearing Celtic brachiopods were geographically closer than had been generally acknowledged. If correct, a new set of plate interactions is needed to explain the evolution of the Appalachian margin. From a South American perspective, much additional geological and palaeontological data from the Palaeozoic rocks of the Andean Cordillera will be necessary to clarify the early history of the pre-Andean margin.

I thank T. M. S/mchez for her ideas during the writing of the paper and D. Harper for helpful comments on the manuscript. I also benefited from comments and discussions with many colleagues, in particular M. Carrera and M. J. Salas. I am specially indebted to R. J. Pankhurst and the reviewers for revision of the English translation. The field work was supported by the Secretarla de Investigaciones Cientificas, C6rdoba University (Project 'The Early Palaeozoic of the Famatina Range') and by grants from Consejo de Investigaciones Cientificas y Tecnol6gicas de la Provincia de C6rdoba (CONICOR) and Consejo de Investigaciones Cientificas y Tecnol6gicas (CONICET). References

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& 1987a. E1 g6nero Platystrophia King (Brachiopoda en la Formaci6n San Juan de la Precordillera Argentina. Ameghiniana, 24, 51-59. -& - - 1 9 8 7 b . Sanjuanella, un nuevo g6nero de la subfamilia Ahtiellinae (Brachiopoda, Plectambonitacea) del Ordovicico de la Precordillera Argentina. Congreso Latinoamericano de Paleontologia, Santa Cruz de la Sierra, Actas, IV, 97-109. & - - 1 9 9 3 . New Early Ordovician Leptellinidae (Brachiopoda) from the San Juan Formation, Argentina. Ameghiniana, 30, 39-58. -& SANCHEZ, T. M. 1979. Modelo de desarrollo del Oc6ano Protoatlfintico en la regi6n Norte de Sudambrica. IV Congreso Latinoamericano de Geologia, Trinidad & Tobago, 2, 825-844. -& 1996a. Paleobiogeography of brachiopod and molluscan faunas along the South American margin during the Ordovician. In: BALDIS, B. & ACElqOLAZA, G. (eds) El Paleozoico inferior del Noroeste de Gondwana. Serie Correlaci6n Geol6gica, Universidad Nacional de Tucumfin, 12, 23-38. & 1996b. The "Afro-South American Realm", and Silurian "Clarkeia Fauna". In: COOPER, P. & JIN, JISUO (eds) Brachiopods, A. A. Balkema, Amsterdam, 29-33. & BRUSSA,E. 1992. Las cuencas silfiricas de Am6rica Latina. In: GUTIERREZ-MARCO, J. C., SAAVEDRA, J. & R.~.BANO, I. (eds) Paleozoico Inferior de Ibero-AmOrica. Universidad de Extremadura, Spain, 119-148. --, SANCHEZ,T. M., CARRERA, M., BRUSSA, E. D. & SALAS, M. J. in press. Paleontological constraints on successive paleogeographic positions of Precordillera terrane during the Early Paleozoic. In: KEPPIE, D. & RAMOS, V. A. (eds),

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The Argentine Precordillera: its odyssey from the Laurentian Ouaehita margin towards the Sierras Pampeanas of Gondwana PATRICIA

WOOD

DICKERSON

1 & MARTIN

KELLER 2

1NASA/Lockheed-Martin, Johnson Space C e n t r e - C23, Houston, Texas 77058, USA 2 Institut fu'r Geologie und Mineralogie, Universitdt Erlangen, Schlossgarten 5, D-91054 Erlangen, Germany Abstract: The Argentine Precordillera (Cuyania terrane) and the Marathon/Solitario basin

(west Texas), both underpinned by Grenvillian Laurentian basement, evolved together during Early Cambrian to Mid-Ordovician times. Ages, Pb-isotope and geochemical data for Precordilleran and for west-central Texas basement (including the Llano uplift) are strikingly similar. Carbonate platform sequences developed on both sides (e.g., E1 Paso and Chica de Zonda) of the Marathon/Solitario outer-shelf to slope basin and hosted homologous reef organisms which were unique to the Laurentian Ouachita margin. The Marathon/Solitario basin received sediments from both north and south; much detritus came from the northern shelf, including olistoliths bearing shelf fauna. Erosional vacuities on the platforms correlate with coarse detrital basin deposits. Cuyania constituted the long-sought southern source. Stratigraphy and structures of the Precordillera/Marathon/Solitario basin are consonant with plate reconstructions that place southern Laurentia near western Gondwana from Late Cambrian into mid-Ordovician time. During Caradoc time, Cuyania moved beyond range of faunal exchange with Laurentia, and tholeiitic basalts were intruded into off-shelf turbidites in the western Precordillera. The attenuated, thermally weakened, Laurentian slab broke apart with continued extension and right-oblique separation of Laurentia and Gondwana. Severance of Cuyania from Laurentia was complete before the onset of Taconic or Ocloyic orogenesis.

A large crustal fragment within the Andes (the Precordillera of western Argentina) has recently been interpreted as a Laurentia-derived terrane (Dalla Salda et al. 1992a, b; Astini et al. 1995, 1996; Thomas & Astini 1996). Ramos (1995) considered that the Precordillera proper is part of a much larger 'Cuyania' Terrane which includes the Grenvillian rocks of the westernmost Sierras Pampeanas (e.g., Sierra de Pie de Palo) and a belt of isolated carbonate outcrops from the province of Mendoza (e.g., San Rafael) southeastwards towards the province of La Pampa (LimayMahuida). Thus defined, the Cuyania terrane would be almost 1000 km long and at least 400km wide, if Andean crustal shortening is restored (Figs 1, 2). The most likely place of origin in Laurentia is the Ouachita embayment (sensu Viele 1989; Dalla Salda et al. 1992a, b) of Arkansas, Oklahoma and Texas, which is distinctly different from the adjacent Appalachian mountain chain in its structural and sedimentary record. The Palaeozoic succession up through lowermost Permian strata of the region was deformed during the late Palaeozoic Ouachita orogeny; to date, however, no Late Ordovician (Taconic) contractional event has been documented (King 1937; Viele & Thomas 1989). The Ouachita embayment has notably similar dimensions to the Cuyania

terrane: the distance between its northeastern boundary, the Oklahoma-Alabama transform, and the southwestern boundary, the Texas transform (Thomas 1991), is of the order of 1000 km. At the southwestern end of the embayment is the Marathon/Solitario area of west Texas, the geochemical, litho- and bio-stratigraphical record of which is entirely complementary to that of the Precordillera (Figs 3, 4). Speculations about the provenance of the Precordillera received new impetus with the discovery of meta-bentonites in mid-Ordovician shelf strata of the Precordillera (Kolata et al. 1994; Huff et al. 1995). The premise originally investigated was that the Precordilleran deposits might be correlative with the well-known metabentonites of the central Appalachians and Baltica (Huffet al. 1992; Haynes 1994). However, the Precordilleran meta-bentonites are older (Arenig-Llanvirn) than those of the Appalachians (Caradoc); Huff et al. (1995) reported a zircon U - P b age of 461+%Ma from the Precordilleran suite, whereas Haynes (1994) cited an age of 453-454 Ma for the southernmost dated Caradoc sample from Alabama. Recently, ashrich subaqueous pyroclastic flows and volcaniclastic debris flows of similar age to that of the Precordilleran meta-bentonites have been found in the westernmost outcrops of the Ouachita fold

DICKERSON,P. W. & KELLER, M. 1998. The Argentine Precordillera: its odyssey from the Laurentian Ouachita margin towards the Sierras Pampeanas of Gondwana. In: PANKHURST,R. J. 8r RAPELA,C. W. (eds) The ProtoAndean Margin of Gondwana. Geological Society, London, Special Publications, 142, 85-105.

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Fig. 1. (a) Geological provinces of western Argentina, showing location of the Precordillera (PR) block: Cordillera de la Costa (CC), Cordillera Frontal (CF), Cordillera Principal (CP), Sierras Pampeanas Occidentales (PO) and Pie de Palo (p), and Famatina (F). (b) Precordillera. Numbered outcrop areas are discussed in text and illustrated on Fig. 7: Guandacol (1), Los Tuneles (2), Portezuelo del Tontal (3), Los Sombreros (4), Cerro Pelado (5), San Isidro (6) (after Astini et al. 1996, figs 1, 3; published with permission, Asociacion Geologica de Argentina).

belt in the Marathon/Solitario area of west Texas (Keller & Dickerson 1996). In the Marathon/ Solitario succession, a notable unconformity spans the time interval during which the younger bentonites were erupted and deposited in Baltica and the central Appalachians (Bergstr6m 1978). Deposits of a two-sided early Palaeozoic sedimentary basin have long been documented in the Marathon/Solitario region (King 1937; Wilson 1954; Young 1970). The Cambrian-Ordovician sedimentary succession is characterized by a variety of siliciclastic and detrital carbonate rocks, which reflect deposition in outer-shelf and upper-slope settings under unstable crustal conditions. Many sediments were derived from the adjacent Laurentian craton to the north; however, felsic volcaniclastic and meta-igneous detritus in the sandstones that could not have been derived from Laurentia (King 1937; Anan 1965; Young 1970; Goter 1973; Calkins 1980)

lends support to the old concept of an enigmatic southern landmass known as 'Llanoria' (e.g., Dumble 1920; Miser 1921; King 1937; Fig. 5). Llanoria has had a chequered past spanning more than 200 Ma, which is discussed in greater detail in Dickerson (in press). Until recently it has been viewed as a single landmass that fed sandto-boulder sized detritus to both the Ouachita and Marathon reaches of the embayment, in both early and late Palaeozoic times (Sellards 1932; King 1937). Llanoria was originally a Carboniferous construct. During Ouachita-Marathon orogenesis, Yucatfin functioned as the late Palaeozoic southern landmass, which encroached and delivered sediment northward to the subsiding Ouachita-Marathon foredeep. The Llanoria concept was subsequently extended back in time to account for a source of non-Laurentian detritus in lower Palaeozoic sedimentary strata in the Ouachita embayment.

THE PRECORDILLERAN ODYSSEY

87

Fig. 2. Northward view of Chile and western Argentina, including the Precordillera. Famatina range (F) and Western Sierras Pampeanas (WSP), southward to near San Rafael. Sierra de Pie de Palo (PdP) in centre. Pacific Ocean (P) at far left; Laguna Llancanelo is the bright area in bottom centre; cities of San Juan (S) and Mendoza (M). (NASA-Mir photograph NM 21-758-085). King (1937) was the first to attribute much of a r61e to Llanoria in the early Palaeozoic; he invoked the landmass as a southerly sediment source for the Cambrian-Ordovician deposits of the Marathon/Solitario basin. Mexico, however, could not have been the early Palaeozoic Llanoria, as it did not arrive on the southern shores of Laurentia until the Devonian (Ruiz et al. 1988; Ortega-Guti6rrez 1994; Keppie & Ortega-Guti6rrez 1995). The presence of Arenig-Llanvirn meta-bentonites in the Precordillera and ash-rich pyroclastic flows in the Marathon/Solitario region lends new support to the ideas of Dalla Salda et al. (1992a, b), Dalziel et al. (1994), and Dalziel (1995, 1997), who claim that Laurentia and western Gondwana were close during Ordovi-

cian time (Fig. 6). In this context, the Cuyania terrane may represent the vanished southeastern continent of Llanoria of early Palaeozoic time.

Cuyania terrane The Cuyania terrane (Fig. 1) consists of two fundamentally different components: (1) metamorphic and igneous basement rocks and (2) the sedimentary cover of Phanerozoic deposits. Those Phanerozoic rocks are the main constituents of the Precordillera of western Argentina. Basement rocks are well exposed in the Sierra de Pie de Palo and related outcrops (Ramos et al. 1993). Their main characteristics are zircon U - P b ages indicative of a major Grenvillian

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Fig. 3. Ouachita embayment of the Laurentian margin, from Alabama westward through west Texas. The Marathon basin, Solitario (S), and Devils River uplift (DRU) are indicated. The city of E1 Paso (EP) occupies the westernmost corner of Texas (Viele 1989).

tectono-magmatic event (McDonough et al. 1993) and a Pb-isotope signature that is typical of the North American (Laurentian) Grenville belt, particularly of Proterozoic basement of the Llano region of Texas (Kay et al. 1996). Deformational history, metamorphic grade, and composition also correspond well with those of the Llano region. The varied lithotypes include meta-rhyolites and meta-andesites (Vujovich et al. 1994). The metamorphic-igneous complex is covered by meta-quartzites and meta-carbo-

nates of the Caucete Group (Borello 1969) or Caucete Metamorphics (Dalla Salda & Varela 1984), which formed the initial sedimentary cover and were metamorphosed during OrdovicianDevonian times. Within the Precordillera, the contact between the basement and the overlying carbonate platform rocks is not exposed. Within the Phanerozoic sedimentary cover of the Precordillera, two different sequences are distinguished: (1) a Lower Cambrian to lower Middle Ordovician carbonate platform and (2) a

THE PRECORDILLERAN ODYSSEY

89

Fig. 4. Northwestward view of the Big Bend region of west Texas. The sharp bend of the Rio Grande (RG) in Big Bend National Park (bbnp) is within the blade-shaped mountain at bottom centre. The Solitario (S) is the circular feature, a breached dome; the Marathon fold-thrust belt (M) is the rectangular area with crenulated rim in right centre. Orientation of the Marathon/Solitario basin in this photo is fortuitously like its early Palaeozoic eastnortheasterly orientation. (NASA Space Shuttle photograph STS 060-85-OB).

Fig. 5. Early Palaeozoic palaeogeography as sketched by King (1937). The Laurentian carbonate platform (c) lay to the northwest and the off-shelf Marathon/Solitario basin (a) was seaward of the platform. The now-vanished landmass of Llanoria, which formed the south flank of the Marathon/Solitario basin, was the southerly source of the abundant igneous, meta-igneous and metasedimentary detritus (b) in early Palaeozoic sedimentary strata. On the basis of geochemical, litho- and biostratigraphical data, the Argentine Precordillera is a likely candidate for that landmass. variety of siliciclastic rocks which are attributed to shallow sub-tidal to deep basin environments. Passive-margin carbonate platform rocks are present from Early C a m b r i a n time o n w a r d , a l t h o u g h the base of a c o n t i n u o u s c a r b o n a t e succession is no older than late M i d - C a m b r i a n

(A. R. Palmer pers. c o m m . 1995); the relation between L o w e r and Middle C a m b r i a n rocks still has to be established. C o n t i n u o u s sections are present a r o u n d the city of San J u a n and near G u a n d a c o l in the n o r t h e r n Precordillera. The following discussion deals mainly with the sections

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Fig. 6. Early to Mid-Ordovician global reconstruction suggesting the proximity of western Argentina and the Ouachita embayment of Laurentia. Baltica (B), East Gondwana (EG), Laurentia (L), West Gondwana (WG). South polar perspective (Dalziel 1995). around San Juan (Sierras de Villicum and Chica de Zonda), which are the classical localities of Precordilleran carbonate studies. Lower Cambrian rocks are not widespread; they occur as tectonically isolated outcrops around San Juan, as an olistolith in the Ordovician slope deposits of the Los Tuneles section, and finally as a redbed-evaporite succession near Guandacol. The Middle Cambrian is represented by the majority of the La Laja Formation (Fig. 7). The main features are thick carbonate units, shoaling upward from sub-tidal environments to probable lower inter-tidal units. They are overlain by relatively thin siliciclastic units, which are interpreted as representing a similar depositional environment to that of the top of the underlying carbonates (Keller & Bordonaro in press). Among the siliciclastic rocks, siltstones are most abundant, although near the base of the sections calcareous sandstones and quartzites are present. Rocks of another Middle Cambrian facies are present in carbonate olistoliths within the Ordovician slope facies (Los Sombreros Fro). Dark grey to black, thin- to medium-bedded limestones are mainly mudstones and wackestones. Some graded beds are interpreted as distal turbidites. Slump features are present within most olistoliths, which may be as thick as 300 m and more than 1 km long (Los Tuneles section). The fauna is dominated by agnostid trilobites and some sponge spicules. The depositional environment of this facies, for which lower and upper contacts as well as lateral relationships are

unknown, is interpreted as an outer shelf to slope setting, well below storm wave base and deep enough to allow a rich agnostid fauna to flourish. Upper Cambrian rocks are mainly dolomites, subdivided into two formations. The Zonda Fm (Fig. 7) consists predominantly of microbial laminites and dolo-mudstones. Oolites, intraclast beds and wackestones are of lesser importance. In contrast to the overlying La Flecha Fm, stromatolites are rare in these deposits. The Zonda Fm was deposited mainly in inter-tidal environments (laminites), although sub-tidal mudstones are also present. The La Flecha Fm is a highly cyclical unit, in which environments from the sub-tidal to the supra-tidal are documented (Armella 1989a, b) and even terrestrial calcretes have been described (Keller et al. 1989). The sub-tidal deposits are characterized by oolites, intraclast conglomerates and thrombolites. In addition, bioturbated mudstones are present. The inter-tidal deposits show a variety of stromatolites (Laterally Linked Hemispheroidal, Solitary Hemispheroidal) and microbial laminites, whereas the supra-tidal rocks consist of microbial laminites and storm-induced mud layers which both show desiccation features and evaporites. Chert, chalcedony, and micro-quartz are very abundant and concentrated in oolite beds and stromatolites. An important feature in the Guandacol section is a lithoclast bed with coarse detrital quartz grains. An interesting Upper Cambrian succession is present at Cerro Pelado (Fig. 7), where rocks like those of the La Flecha Fm (Cerro Pelado Fm of Heredia 1990) are overlain by deep-water, agnostid-bearing limestones that grade up-section into predominantly marlstones and shales (El Relincho Fm of Heredia 1990). The contact between the units represents a drowning unconformity (sensu Schlager 1989). A similar event has not yet been documented in the carbonate platform succession proper. Finally, deep-water carbonate olistoliths like those of the Middle Cambrian are present in the Los Sombreros Fm. The Early Ordovician platform succession starts with medium- to thick-bedded limestones (La Silla Fm; Tremadoc) and minor intercalations of microbial laminites and thrombolites. Dominant lithotypes are mudstones and wackestones, the latter containing principally gastropods and nautiloids, and peloidal grainstones and intraclast grainstones. During the Arenig (San Juan Fm) open-marine carbonates prevailed, with an abundant and diverse fauna. Most important are reef-mound zones near the base and near the top of the formation. The lower mounds are composed of sponge-algalreceptacutitid communities, whereas the upper

THE PRECORDILLERAN ODYSSEY

91

Fig. 7. Stratigraphical columns for the Precordillera/Marathon/Solitario basins. See Figs ! and 3 for locations of sections.

reefs are dominated by stromatoporoids with minor amounts of sponges and algae. The upper San Juan Fm contains several meta-bentonite beds (Kolata et al. 1994; Huff et al. 1995). Their intercalation with well-dated fossiliferous Llanvirn strata makes them useful for intercontinental correlation; the oldest bentonites so far recorded are of middle Arenig age (Bergstr6m pers. comm. 1994). During the earliest Llanvirn, carbonate platform evolution was terminated by deposition of graptolitic black shales (Gualcamayo Fm), which cover the carbonates above a drowning unconformity. Near Guandacol, in the northernmost part of the Precordillera, this event was earlier and dated as mid-Arenig. Finally, Lower Ordovician rocks are present within the Los Tuneles olistolith, where Tremadoc debris flows and turbidites erosionally overlie Middle Cambrian agnostid-bearing limestones (Keller 1995).

About 25 m of slope deposits include Tremadoc, Arenig and earliest Llanvirn faunas. Lower Ordovician rocks are also present south of San Rafael at Ponon Trehue (Bordonaro et al. 1996; Keller 1996). There, about 80m of dolomites and limestones represent much of the Tremadoc to middle Arenig record. These strata, which rest on Grenvillian basement (Ramos pers. comm. 1996), are mainly granitoids, gneisses and some meta-siliciclastic rocks. The overlying sediments can be correlated with the Lower Ordovician type sections in the Precordillera at Cerro La Silla. It can be demonstrated that about 100 m of carbonates at Cerro La Silla correspond to about 20 m of the Ponon Trehue Fm. Nevertheless, all major sedimentological elements, including the lower reef mounds, are developed. The Mid- and Late Ordovician were times of major facies differentiation. After the regional

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P. W. DICKERSON & M. KELLER

drowning of the platform, shallow-marine carbonate sedimentation continued only locally. The Las Aguaditas Fm (Middle to lower Upper Ordovician) is composed of deep-water limestones and marlstones with abundant intercalations of sediment-gravity flows (Keller et al. 1993). The depositional environment is interpreted as a slope apron developed along normal faults. Sediment transport was apparently from east to west. Along the western margin of the former platform a thick succession of continental-slope sediments was deposited (Los Sombreros and Empozada Formations, Fig. 7). The main characteristics of both units are matrices of shales and marlstones which host sediment gravity flows. A unique association in some of the olistostromes are huge Middle Cambrian olistoliths together with basement boulders (Banchig et al. 1990; Keller 1995). The whole platform succession is represented by the clasts within the turbidites and mega-breccias. The formation of olistoliths as thick as 300m and more than l k m long requires steep escarpments, along which the basement of the Precordillera may locally be exposed (Keller 1995). Boulders smaller than the mega-olistoliths can be transported on shalelubricated slopes of 5 ~ to 10~ as documented by Cook (1970) in the Aleutian Islands. Coeval slope and basin sediments include the Portezuelo del Tontal Fm, the clast content of which is similar to that of the Los Sombreros and Empozada Formations and which also indicates a source area to the east (Spaletti et al. 1989). Although the relations among the various Middle and Upper Ordovician basin deposits are not clear, the general tendency is that of fining toward the west. In this context, the Portozuelo del Tontal Fm. (Figs 1, 5) is geographically near the Los Sombreros Fm and probably reflects more distal slope development.

Marathon[Solitario basin, Ouachita embayment Grenvillian igneous and metamorphic basement rocks ranging in age from 1 to 1.3 Ga form the underpinnings of this span of the Laurentian margin (Mosher 1993). Common-lead isotope data indicate that Laurentian crust extends as far south as the Marathon/Solitario region and that it was modified during late Palaeozoic contractional orogenesis (James & Henry 1993). Seismic reflection, potential fields, and field structural data demonstrate that basement rocks are block-faulted (generally down to the south-

east) and that pre-existing fault-block topography exerted marked control on late Palaeozoic fold/thrust geometry (Tauvers 1988; Muehlberger 1995). NE-trending fault-bounded blocks also influenced early Palaeozoic sedimentation (Anan 1965; Young 1969, 1970; Goter 1973; Calkins 1980; Canter 1993). Interpretation of seismic reflection and well data undertaken while the first author of this paper was affiliated with Gulf Research & Development Co. (c. 1980), with Conoco (1985), and later as a consulting geologist for a consortium of petroleum exploration companies ('Controls on Reservoir Development and Distribution in the Ellenburger Group, Val Verde Basin', unpublished report by Coyote Geologic Services 1991), corroborates the cited field observations of fault control on early Palaeozoic sedimentation and on late Palaeozoic fold-thrust geometry. Cambrian to Middle Ordovician rock types for both shelf and off-shelf sequences are broadly comparable, although modes of deposition and fauna differ (Dickerson 1993). At the base is a time-transgressive siliciclastic unit of highly variable thickness. Lower Ordovician lithic suites are mainly limestone and dolomitized limestone, with subordinate calcareous shale and sandstone. Following a hiatus of variable duration, depending upon palaeo-topographical position, voluminous Middle Ordovician conglomerates, sandstones (including calcarenites) and shales were deposited. Much detritus in the off-shelf sediments consists of material derived from nearby shelf strata including large olistoliths bearing shelf fauna. Erosional vacuities on the shelf correlate with coarse detrital deposits in basin strata. A significant difference, however, between the shelf and off-shelf suites is the abundance in offshelf deposits of clasts from a mixed volcanic/ sedimentary/metamorphic source to the south (King 1937; Wilson 1954; Young 1970). Compositions of the commonest felsic volcanic clasts (trachyte, dacite, trachyandesite, felsic porphyry) require a source other than the Devils River uplift (Fig. 3), where the only known southern Laurentian Ordovician volcanic strata (rhyolite) have been analysed (Nicholas & Rozendal 1975). Cambrian volcanic and sedimentary strata of the southern Oklahoma aulacogen (granite/rhyolite, gabbro, unmetamorphosed greywacke) are not of the requisite composition (Muehlberger et al. 1967). Furthermore, neither they nor the Precambrian-Cambrian metasedimentary rocks in northeastern New Mexico and the Texas Panhandle were geographically positioned to satisfy the observed variations in thickness (northward thinning) and abundance (northward decrease).

THE PRECORDILLERAN ODYSSEY The Upper Cambrian (Trempeleauan lower Tremadoc) Dagger Flat Sandstone (Fig. 7) is the oldest unit exposed in the area (King 1937; Wilson 1954). Sands and shales of the formation, which contain metre-scale olistoliths, were deposited in tectonically unstable outer-shelf to upper-slope environments. Sources for the sands lay both to the south and to the north. The northern source supplied metamorphic and granitic clasts; metamorphic rock fragments amount to as much as 42% of Dagger Flat sands in places. Granitic and meta-igneous rocks were exposed along the margin immediately north of the Marathon basin at that time. From the south came volcanic rock fragments (up to 20% of some samples), basic igneous rocks (Anan 1965; no compositional details), and green primary chert clasts. Sporadic increases in volcanic rock fragments from a southerly source could be '... explained by volcanic activity accompanied by some structural activity, such as block-faulting' (Anan 1965). Block faulting has also been proposed by Calkins (1980) as an influence on Dagger Flat deposition farther west in the Solitario. Anan (1965) suggested an array of NEtrending, fault-bounded, high and low blocks to account for lithofacies variations within the Dagger Flat Fm. Concentrations of clean quartz sandstones and flat-pebble conglomerates could reflect shoaling on topographically higher blocks. Older Cambrian shallower-water limestones preserved on downfaulted blocks might also furnish local sources of such boulders which are found in the much younger Woods Hollow Shale (Llandeilo-Caradoc). The disparate interpretations of shallow versus deep marine depositional settings for almost all the Lower Palaeozoic strata of the Marathon succession might be reconciled within an outer-shelf setting of highly variable basin-floor topography. Dagger Flat sandstones and shales grade upward into the shales of the Lower to Middle Ordovician Marathon Fm (Tremadoc-Arenig; Wilson 1954). The upper one-half to one-third of the formation is thin-bedded fossiliferous detrital limestone. Young (1970) reported abundant volcaniclastic material (rhyolite, dacite, trachyte, trachyandesite, some andesite) in the sandier intervals of the Marathon Fm and stated that the volcanic debris came from a southerly source. Dacite, trachyte and trachyandesite sources of appropriate age are unknown in west Texas, New Mexico and Oklahoma, and these lithoclasts become progressively more abundant southward. Thick, mature to supermature quartz arenites are found high in the Marathon Fm in the Solitario and in the south-

93

ern reaches of the Marathon basin but not in the north. Until the recent recognition of ash-rich subaqueous pyroclastic flows in the Fort Pefia Fm (Keller & Dickerson 1996), there was only one mention of bentonite in any of the Lower Palaeozoic strata of the region, that having been from drill cuttings; C. L. Baker & D. O. Carsey (in King 1937) described bentonite interbedded with bluish-purple chert in the upper Marathon limestone that had been penetrated in a well near the town of Marathon. Keller & Dickerson (1996) reported the presence of a 3 cm thick bentonitic bed within the upper Marathon limestone in the Solitario, establishing that the effects of late Arenig explosive volcanism were felt in this region, as they were in the Precordillera (San Juan Fm). Volcanic origin of the Solitario occurrence was established primarily on the basis of its euhedral zircon, 3-quartz, and apatite content. Olistostromes both low and high in the formation attest to unstable conditions during its deposition. Olistoliths of coarse sandstone like that of the Dagger Flat are present in the lower portions, whereas dolomitized carbonateplatform limestone boulders that constitute the Monument Spring Member are found at higher levels. The carbonate boulders appear to have been derived from local fault blocks in the northwestern basin (King 1937; Wilson 1954; Young 1969, 1970; McBride 1989); they are not found farther south in the Marathon region or in the Solitario. There is a minor hiatus of variable duration between the Marathon Fm and the Alsate Shale/ Rodriguez Tank Sandstone (Izold 1993), as there is between the San Juan Fm and younger units in the Precordillera (Baldis et al. 1985). The Alsate Shale of the Marathon basin and the laterally equivalent Rodriguez Tank Sandstone at the type locality in the Solitario are lower Llanvirn. The Alsate consists largely of dark grey graptolitic shale with lesser limestone (graptolite zone 8 - lsograptus caduceus of Berry 1960); limestone becomes more abundant to the south (King 1937). Quartz sand is markedly more abundant in the southern Marathon basin and in the Solitario, at the level of the Alsate/ Rodriguez Tank formations and within the overlying Fort Pefia Fm. Detailed mapping indicates that not all the sands previously mapped as Rodriguez Tank in the Solitario are Alsate-equivalent; several have now been recognized as being within the overlying Fort Pefia Fm (Muehlberger 1996). The Middle Ordovician Fort Pefia Fm (Llanvirn lower Llandeilo) conformably overlies the

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P. W. DICKERSON & M. KELLER

Alsate and Rodriquez Tank formations (Berry 1960; Goter 1973) and consists of calcarenites, conglomerates, both rounded and flat-pebble, abundant bentonite-derived shale (Weaver 1961), and siliciclastic beds that are both normally and reversely graded. Interbedded with those strata are ash-rich subaqueous pyroclastic flows and volcaniclastic debris flows, discussed in detail below ('Arenig-Llanvirn Volcano-sedimentary Deposits'). All are interpreted as having been deposited in an outer-shelf setting, at or somewhat below normal wavebase (Goter 1973). Isotropic silica is preserved in rare samples from the Solitario and is probably from an igneous source (Calkins 1980); numerous reddish to purplish chert beds, commonly indicative of relict volcanic ash layers (Conkin & Conkin 1983), are characteristic of the formation in the Marathon basin and in the Solitario (King 1937; McBride 1989; Calkins 1980; W. R. Muehlberger pers. comm. 1996). Phosphatic shale and glauconite are common as well (Goter 1973). There are boulder beds with clasts up to 6 m long in the lower part of the Fort Pefia Fm in the Marathon basin (Goter 1973; McBride 1989), but none has yet been noted in the Solitario. The

olistostromes occupy a position roughly equivalent to those of the Los Sombreros Fm of the Precordillera, reflecting unstable depositional conditions and probable syn-depositional faulting in both areas during Llanvirn-Llandeilo time. Boulders consist principally of shallowmarine limestone with subordinate arkosic sandstone; andesite and granite clasts are also present. The Fort Pefia Fm grades upward into the Woods Hollow Shale (Llandeilo-lower Caradoc). The Woods Hollow Fm consists dominantly of silty phosphatic clay shale with subordinate calcareous fine sandstone and limestone (McBride 1989); the clays are bentonite-derived mixed-layer illite/smectite [montmorillonite] (Weaver 1961). Current-rippled beds, flat-pebble conglomerate zones, and other sedimentary structures observed in the Woods Hollow Fm in both the Marathon basin and the Solitario indicate deposition of much of the unit in waters at or above storm-wave base, possibly on a flanking beach of a volcanic edifice (Lajoie & Stix 1992; Fig. 8). The Woods Hollow Fm contains large olistoliths similar to those in coeval Los Azules and La Cantera formations of the Precordillera;

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/

Fig. 8. Idealized vertical and lateral facies variations in subaerial and subaqueous environments for a marine volcanic vent. Normally graded, cross-bedded and stratified facies, massive and reversely graded facies, and deposits interpreted as remobilized shoal or possibly beach sediments are represented in the Fort Pefia Fm in the Solitario (from Lajoie & Stix 1992; published with permission, Geological Association of Canada). Subaqueous pyroclastic flows have been observed to retain coherence to distances >250 km from the source vent (Sparks et al. 1980a).

THE PRECORDILLERAN ODYSSEY included are platform limestone blocks with faunas ranging in age from early Late Cambrian (possibly latest mid-Cambrian) to Early Ordovician (late Tremadoc; Wilson 1954). Late Cambrian faunas in the Woods Hollow boulders are older than those in any Cambrian strata now exposed in the Marathon/Solitario region; given the block-faulting that persisted throughout Late Cambrian to mid-Ordovician sedimentation, older Cambrian rocks could well be preserved on down-dropped blocks within the basin. Sandstone, schist and felsic igneous blocks are also present and believed to have come from a southerly source; Devils River uplift could have provided clasts of those compositions but the areal extent and geometry of those strata are unknown. In the outer-shelf to upper-slope sequence of the Marathon/Solitario region there is a significant unconformity at the top of the lowermost Caradoc Woods Hollow Fm/base of the Ashgill Maravillas Fm (Baker & Bowman 1917; King 1937; Wilson 1954; Bergstr6m 1978). On topographically high blocks within the platform sequence the entire mid-Ordovician section is missing at the unconformity; in the Franklin Mountains near E1 Paso, for example, the break spans c. 30 million years (Fig. 9). The unconformity at the top of the El Paso Group is the post-Sauk unconformity recorded in Laurentian platform carbonate suites from Newfoundland to at least eastern New Mexico. In west Texas, Lower Ordovician carbonate strata were blockfaulted and subjected to pervasive karstification prior to deposition of the Ashgill Montoya Fm. In topographically lower areas of the platform, mid-Ordovician shale and sandstone of the Simpson Group fill both fault and karst topography on the Lower Ordovician limestones (Galley 1958; Dyer 1989; Canter 1993); Montoya Fm carbonates were then deposited upon the Simpson siliciclastic strata. In the world beyond the Ouachita embayment, development of the Caradoc unconformity began at the same time that volcanism ceased in the Famatina system. Significantly, inherited zircons from Famatinian rocks yield a protolith age of 1.7Ga (Pankhurst et al. 1996, this volume), which indicates that the continental component there was not the Grenvillian Laurentian continental crust of Cuyania. In addition, the break spans the time of eruption and deposition of the Caradoc meta-bentonites of Appalachia and of Baltica (Kolata et al. 1994; Haynes 1994; Huff et al. 1995). Emplacement of pillow basalts in the lower Caradoc Alcaparrosa Fm of the western Precordillera indicates that separation of that area from Laurentia was then

95

Fig. 9. Post-Sauk unconformity in the Ordovician platform carbonate succession of the Franklin Mountains, El Paso, Texas. The height of the cliff face from road level (top of steps, lower right) to top of dark grey unit is c. 35 m. Darker dolomite and chert of the Montoya Group (Ashgill) directly overlie the lighter limestones and subordinate sandstones of the E1 Paso Group (Arenig). Llanvirn to early Ashgill deposits are absent.

well under way. As discussed in detail by Benedetto and Keller et al. (both in this volume) the Precordillera finally moved beyond the range of faunal exchange with Laurentia during early Caradoc time, thus ending the shared history of the two areas.

Arenig-Llanvirn

volcanogenic deposits

The discovery of meta-bentonites in ArenigLlanvirn shelf-carbonate strata in Cerro Viejo and Sierra de Talacasto of the Precordillera (Bergstr6m et al. 1994; Huffet al. 1995) permitted a first-order test of reconstructions placing the Precordillera and Marathon/Solitario region in proximity at that time (Dalla Salda et al.

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1992a, b; Dalziel et al. 1994; Dalziel 1995). The Fort Pefia and Marathon formations were selected for examination based on biostratigraphical information from S. Bergstr6m (pers. comm. 1994). Those formations are correlative with meta-bentonite-bearing carbonateshelf strata of the Precordillera, so presence or absence of meta-bentonites could support or refute proximity of the two regions. The presence of ash-rich subaqueous pyroclastic flows and bentonitic debris flows in the Fort Pefia Fm in the Solitario was established in 1996 (Keller & Dickerson 1996). Later that year Keller (in a field party with R. Astini, L. Dalla Salda, I. Dalziel, P. Dickerson, F. Hutson, S. Kay, V. Ramos, & W. Muehlberger) found a 3 cm thick bentonitic bed in the Marathon Fm in the Solitario (Keller & Dickerson 1996). The volcanic origin of that layer was established on the basis of its crystalline constituents: euhedral zircon,/3-quartz and apatite. We regret that the term 'meta-bentonite' was used imprecisely in earlier reports, as the bentonitic material is mixed with normal marine sediment in this outer-shelf to slope setting. We here summarize what is known to date about the pyroclastic deposits of the Marathon/Solitario basin, particularly with regard to questions about their identification (Huff, in Dalziel 1997; Huff et al. this volume).

Previously known occurrences. Prior to this work, known Ordovician volcanic rocks in the Marathon/Solitario region included the bentonite bed in the upper Marathon Fm (Arenig) encountered in a borehole near the town of Marathon. A slightly metamorphosed Ordovician felsic volcanic rock dated at 481 • 20 Ma (Rb-Sr; Nicholas & Rozendal 1975) was penetrated in a borehole on the Devils River uplift. Significantly, with respect to a source for igneous/meta-igneous clasts in Lower Ordovician sedimentary strata, no dacites, trachytes, trachyandesites or andesites were present. King (1980, p. 37) describes andesite intrusive rocks in mid-Ordovician strata in a borehole 3 km SE of the town of Marathon. Stratigraphical assignments differ for the section in sample logs for the well (Marathon v. Fort Pefia); King states that Marathon Fm overlies the Palaeozoic sole thrust (at 4850 or 5980-6100ft) and the intrusive rocks are between 3870 and 3990 ft. The andesites are found above the late Palaeozoic sole thrust and not in the foreland sediments below. The only other known andesites in the region (Campus Andesite, c. 48Ma; Henry & McDowell 1986) are around El Paso and in

sulphur exploration boreholes in the Delaware Basin (G. Foose pets. comm. 1986). In addition, there are numerous intriguing descriptions of green clays, blue to purple cherts, and granular cherts in Lower Ordovician shelf strata of west Texas, which we have not yet examined in the field. Fort Pet~a outcrop. The sequence of yellowishbeige to bluish-grey bentonitic deposits mixed with normal rnarine sediments near the top of the formation in the Solitario is more than 7 m thick and comprises at least 22 event beds; the base of the unit at this locality is a fault contact. The crystalline component, which includes euhedral .%quartz, zircon, apatite, sphene, and minor biotite (commonly altered to chlorite), constitutes up to 80% of some Fort Pefia samples, compared with 10% of the Sierra de Talacasto sample collected and analysed by Dickerson for comparison. The subaqueous pyroclastic flows and volcaniclastic-rich debris flows are at an equivalent to somewhat higher stratigraphical level than are the thin meta-bentonites of the Precordilleran carbonate shelf sequence (S. Bergstr6m pets. comm. 1996); not surprising, given the greater likelihood of wave, storm and faunal reworking on the shelf, as observed by Hanson (1995) and Hanson et al. (1996). Depositional setting. The thickness and crystal content of the Fort Pefia samples, the outer-shelf to slope depositional setting above Carbonate Compensation Depth, the admixture of normal marine carbonate and siliciclastic sediments, and sedimentary structures of varying scales are consistent with deposition as part of the volcaniclastic apron of a submarine to subaerial vent complex. Excellent analogues are illustrated by Lajoie & Stix (1992) and are amply documented in modern volcanic aprons of the Lesser Antilles arc (Carey & Sigurdsson 1984; Reid et al. 1996) and of the New Hebrides arc (Reid et al. 1994), among others. Of particular relevance with respect to the Fort Pefia debris flows is the research of Sparks et al. (1980a, b) on- and offshore Dominica in the Lesser Antilles. They document individual andesitic to dacitic blockand-ash flows as thick as 40 m, which are rich in crystals, poor in shards, entrain boulders up to 10m in diameter, and extend from subaerial to submarine reaches of the vent flanks. Those flows maintained coherence out to 250 km from the source. Composition and geochemist~T. Euhedral /3quartz, zircon, sphene, and apatite of volcanic origin have been separated from all the Fort Pefia samples. Euhedral biotite is also present

THE PRECORDILLERAN ODYSSEY but commonly altered to chlorite. All Fort Pefia samples contained abundant calcium carbonate from marine silt and calcite cements. The crystalline fraction not only establishes the volcanic contribution to the Fort Pefia sediments but should also provide for U-Pb dating (zircon) to refine the graptolite biostratigraphical age. If an analytically significant number of ~-quartz crystals encasing glass inclusions can be extracted, both an accurate age and the composition of the source magma will be determinable. The potential value of glass inclusions in /#quartz is that they can provide vital data on the composition of the melt that was trapped during phenocryst growth within the magma chamber prior to eruption (Delano et al. 1994; Hanson et al. 1996), i.e., under virtually closed-system conditions. Alteration of the bulk composition of a volcanic ash begins the moment it is blown from the vent, starting with physical sorting in air or water, according to its density and aero- or hydro-dynamic properties. In the aqueous environment diagenesis, mechanical and faunal remobilization, and mixing with ambient sediment further modify the composition. Compaction, de-watering, and burial diagenesis contribute. At no stage of meta-bentonite formation do closed-system geochemical conditions prevail. For interpretations of magma-chamber processes, for ash-to-ash correlation and geochemical fingerprinting, and for speculations on tectonic setting for volcanism, chemical data from melt inclusions obviate much of the uncertainty and data manipulation required when meta-bentonite bulk composition is used as a surrogate for melt composition (e.g., Huff et al. 1995, 1996; Bergstr6m et al. 1996; further discussion in Hanson et al. 1996). Conditions of closed-system crystallization of granites, for which the tectonic discrimination diagrams of Pearce et al. (1984) were devised, might be more nearly satisfied when melt chemistry derived from glass inclusions in phenocrysts is used. Age. The Fort Pefia beds contain Llanvirn graptolites assigned by Berry (pers. comm. May 1996) to his zone 9, Hallograptus etheridgei of the Marathon succession. Because the later Arenig to Llanvirn was a time of maximum provincialism (Finney & C h e n Xu 1990), isotopic ages for Marathon graptolite zone 9, when determined, will have broader significance for global correlations of Llanvirn strata. Clay mineralogy. Clay mineralogy is neither a unique discriminant for volcanic origin of clays nor a correlation tool (e.g., Conkin & Conkin 1983; Haynes 1994; L. Lynch pers. comm. 1996).

97

Nonetheless, the compositional data are instructive with regard to the diagenetic history of the volcano-sedimentary strata. X-ray diffraction analysis has been carried out on samples from three ash-rich pyroclastic flows from the Fort Pefia Fm in the Solitario and on one sample collected from the thickest meta-bentonite of the San Juan Fm in the Sa. de Talacasto section of the Precordillera. Like those of the San Juan sample, clays of the Fort Pefia are dominantly illite; diffractograms for all three exhibit distinct illite/smectite (~montmorillonite) peaks, whereas the San Juan data showed none. Chlorite is present in two of the three Fort Pefia clays (the peaks persisted after heating the sample to 400~ and to 550~C). Chlorite was observed in thin section as well, in several instances replacing biotite. Alteration of biotite to chlorite commonly occurs during burial diagenesis (e.g., Haynes 1994); in addition, chlorite is a product of weathering of andesitic volcanic rocks (R. L. Folk pets. comm. 1996). From the mineralogical and sedimentological data so far available, we interpret the Fort Pefia Fm volcanogenic sediments as ash-rich pyroclastic flows and debris flows, deposited in an outer-shelf to slope environment. They, and the bentonitic bed in the Marathon Limestone, may be correlated with, or are slightly younger than, meta-bentonite beds in the shelf-carbonate strata of the Precordillera. The Marathon/Solitario volcanogenic sediments may therefore support reconstructions juxtaposing the two regions in Arenig-Llanvirn times.

Co-evolution of the Precordillera] Marathon/Solitario Basin The Marathon/Solitario Basin was a two-sided basin, not merely an embayment in the passive margin. To the north, it was bounded by the circum-Laurentian carbonate platform, from which it received much carbonate detritus. However, a candidate for the southern counterpart that served as a source for much of the siliciclastic sediment has only recently been recognized (Keller & Dickerson 1996). The Cuyania carbonate platform shows a welldeveloped passive-margin succession. Isolated coeval slope sediments do not appear before the Early Ordovician, and regional slope development did not start before the latest Llanvirn (Keller 1995). Lower Ordovician basin deposits are unknown in the Precordillera. The abrupt facies changes throughout the Marathon/Solitario basin sequence, especially the presence of Upper Cambrian carbonate

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P. W. DICKERSON & M. KELLER

boulders in the Middle Ordovician Woods Hollow Shale, point to considerable syn-sedimentary extension and block faulting of the platform succession. Similarly, the western margin of the Precordillera is characterized by extensional faults from Late Cambrian time onward (e.g., Cerro Pelado section). An important clue to the original platetectonic configuration is provided by the metabentonites (Bergstr6m et al. 1994, 1996; Huff et al. 1995) and ash-rich pyroclastic flows (Keller & Dickerson 1996). In both regions, they are found in the same stratigraphical interval (middle to upper Arenig through to Llanvirn); those of the Solitario persist a little higher into the Llanvirn. Ongoing analyses of the crystalline mineral suite from ash-rich deposits of the Fort Pefia Fm of the Solitario (/3-quartz, apatite, zircon, titanite) will provide meaningful data for testing the correlation of Marathon/Solitario volcanogenic materials in off-shelf strata with those of the Precordillera shelf sedimentary sequence, as well as with the Ordovician rhyolite of the Devils River uplift (Fig. 3). It is noteworthy that both suites of meta-bentonites and ash-rich sediments belong to a distinctly older population (by c. 8Ma) than do those of the central Appalachians and of Baltica (Bergstr6m et al. 1994; Haynes 1994). If the Cuyania terrane and the Marathon/ Solitario area did indeed share a common history from latest Precambrian into mid-Ordovician time, that history would have begun with latest Precambrian to Cambrian rifting, which initiated the opening of an ocean along the eastern Appalachian and southern Ouachita margins of Laurentia. As noted by Thomas (1991), the expression of that margin varies along its length. In contrast with the Appalachian margin, riftedmargin prisms seem to be absent from the Ouachita margin (Viele & Thomas 1989), and they are most probably also absent from the Precordilleran margin (Keller 1995). Their absence is consistent with transtensional separation of the latter two areas as, for example, in the Sea of Japan (e.g., Kong & L a w v e r 1995). Following that, a sedimentary basin developed between the Laurentian margin to the N/NW and the Cuyania continental margin to the S/SE. This basin was flanked on either side by vast carbonate platforms typical of passive margins (e.g., Bond et al. 1984; Gonzfilez Bonorino & Gonzfilez Bonorino 1991). From Mid-Cambrian time onward a deeper water environment persisted on the Precordilleran margin. Disintegration of the carbonate platform along its western margin started during the Late Cambrian (Cerro Pelado section).

The submersion of peri-tidal carbonates beneath deeper water limestones and marlstones, an event which is not observed on the platform itself, points to a sudden foundering of the platform edge, probably along normal faults. The Upper Cambrian succession of the Marathon Basin consists primarily of sandstone. Anan (1965) suggested an array of NE-trending fault blocks to account for lithofacies variations within the Dagger Flat Fm. Boulder beds within the formation attest to unstable conditions during deposition. Onlap of Lower Ordovician carbonates onto exposed basement near San Rafael demonstrates that a source for siliciclastic input into the basin was present during much of the Cambrian and Early Ordovician history. Igneous and metamorphic basement rocks there, as well as in Pie de Palo (Caucete Group), may have provided the exotic siliciclastic detritus to the Dagger Flat, Marathon, Rodriquez Tank and Fort Pefia sandstones in the Marathon/Solitario area. During the Early Ordovician, carbonate sedimentation continued on both platforms and slowly onlapped cratonic basement. In the Cuyania terrane this is documented for the Ponon Trehue section (Bordonaro et al. 1996); on the Laurentian margin an analogous section is the E1 Paso Group near the city of E1 Paso, Texas (LeMone 1988a). Significantly, both platforms

BB continental interior shelf environment

sponge-algal facies

off-shelf Ouaehita facies

hypothetical continuation of carbonate platform, Argentine PrecordiHera

Fig. 10. Distribution of Arenig sponge algal bioherms that rimmed the Ouachita embayment of the Laurentian margin and are found in counterpart strata of the San Juan Fm of the Argentine Precordillera. The star marks the location of E1 Paso, westernmost Texas. (Modified from Alberstadt & Repetski 1989).

THE PRECORDILLERAN ODYSSEY

Marathon-Solitario basin

Laurentian margin

99

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El Paso

eastern basin Zondat

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carbonate platform succession

ocean floor with basinel shales end pillow basalts

syn-rift deposits

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youngest platform deposits and correlatives

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Fig. 11. Palaeogeographical restorations of the Argentine Precordiltera and Marathon/Solitario region of the Ouachita embayment of Laurentia: (a) Early Ordovician, (b) Mid-Late Ordovician. share nearly identical reef structures. In the Franklin Mountains of west Texas the dominant reef organism is Pulchrilamina, a stromatoporoid-like organism (LeMons 1988b), whereas in the Precordillera Zondarella communis is the main organism, also attributed to the stromatoporoids (Keller & Flfigel 1996). The sponge-algal facies that hosts these reefs is mainly distributed around the Ouachita margin of Laurentia (Alberstadt & Repetski 1989; Figs 10, lla). These facies have their clear counterpart in the San Juan Fm, which mimics the Laurentian occurrences even in details like the presence of Nuia and the absence of the Epiphyton/Renalcis community (Cafias & Keller 1993; Keller et al. 1993). Adjacent to the platforms, steep slopes developed which generated sediment gravity flows. Today evidence of them is present within the mega-olistolith in the Los Tuneles section of the Los Sombreros Fro. (Keller 1995) and in the Marathon Fm (McBride 1989).

Carbonate boulders within the Marathon Limestone were locally derived, having come from the block-faulted shelf sequence to the northwest (Wilson 1954; McBride 1989). However, siliciclastic/volcaniclastic sands within the Marathon Fm came from the southeast to southwest (Young 1970) and, as basement rocks in the Cuyania terrane were still exposed in Early Ordovician times, might have been derived from this source, The Arenig-Llanvirn Famatina eruptive complex (Rapela eta/. 1992) could also have constituted a southerly source for fine-grained rhyolitic/dacitic/andesitic detritus in the Marathon and Fort Pefia formations. During Mid- and Late Ordovician times the Precordillera was characterized by large-scale extension. This is documented by the emplacement during Caradoc time of rift-related pillowbasalts within graptolite-bearing (Nemagraptus gracilis zone; Llandeilo) turbidites (Kay etal. 1984) and the formation of horsts and grabens

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P. W. DICKERSON & M. KELLER

in the former platform area (Fig. llb). Largescale extension is also responsible for the formation of steep fault scarps, the source for the mega-olistoliths and the basement boulders in the Los Sombreros Fm (Keller 1995). The stratigraphically equivalent Fort Pefia Fm and the younger Woods Hollow Shale also contain abundant olistostromes with platform carbonate and exotic clasts (schist, sandstone, felsic porphyry; McBride 1989). The presence of Late Cambrian to Early Ordovician fossils in the carbonate blocks points to considerable uplift in the hinterland. Extension and the block-faulted topography in the Precordillera at that time provides a source for the igneous and metamorphic boulders observed in the Woods Hollow Shale. The Arenig-Llanvirn meta-bentonites described from the Precordillera (Kolata et al. 1994; Huff et al. 1995) and volcanogenic flows from the Marathon/Solitario area (Keller & Dickerson 1996) are also interpreted as related to crustal extension, which was the dominant deformational regime during Cambrian to midOrdovician times in the Precordillera (von Gosen 1992; Loske 1992; Keller 1995; Buggisch 1996), as well as in the platform sequence and in the Marathon/Solitario region of west Texas (Young 1969, 1970; Ammon 1981; Dickerson 1980, 1994, 1995; Dyer 1989). The extensional events reflect ongoing separation of the Cuyania terrane from Laurentia, culminating during the late Llandeilo and Caradoc, which is in agreement with the faunal data and with models presented by Dalla Salda et al. (1992a, b), Dalziel et al. (1994), and Dalziel (1995). Keller et al. (this volume) discuss the biostratigraphical data in detail. Sediment input into the western basin was markedly diminished during the Caradoc; in the Precordillera to the east, sediment was trapped in the platform interior grabens and along the continental slope/continental rise. As a result, the Marathon/Solitario basin was converted to a deep-water oceanic basin shut off from detrital sedimentary input, contributing to the lengthy Caradoc hiatus. Several observations must be reconciled in any plate tectonic scheme proposed to account for rifting and transfer of the Cuyania terrane from Laurentia to Gondwana. Based upon the palaeomagnetic reconstructions of Dalziel (1995), for example, during the period from 550 to 420 Ma (particularly between c. 485 and 420 Ma), Laurentia was moving steadily northward and clockwise with respect to Gondwana. During the same period Gondwana rotated more than 90 ~ anticlockwise. Such plate trajectories would require

that the mode of deformation of the basin were right transtension rather than near-orthogonal extension. Ramos' (1984) interpretation of dextral offset of elements of the Gondwanan margin would be consonant with that mechanism and timing for deformation. Separation would have been more complex than rifting away of the transform-fault bounded terrane and translating it eastward across the widening sea (Astini et al. 1995; Thomas 1996; Thomas & Astini 1996). Such rifting and transform faulting in the marine realm would result in generation of oceanic crust, of which only small occurrences have been located (e.g., western Precordillera; Kay et al. 1984); none has yet been established on the Ouachita margin, although Thomas (pers. comm. 1997) speculates that serpentinite pods in the Ouachita thrust belt represent fragments of obducted Iapetus crust. Whatever their age, the bodies are not extensive. Plate rotations of the sort described (Dalziel 1995) are particularly problematical for maintaining the integrity of a Laurentian-basement cored plateau spanning the mid-Ordovician ocean (c. 20 ~ of latitude in the 465 Ma reconstructions of Dalziel 1996, 1997). Regarding incorporation of Cuyania into Gondwana, eastward subduction and corresponding magmatism are well documented along the western Gondwanan margin in the Western Sierras Pampeanas (Ramos 1988, 1995; Vujovich et al. 1994; Astini el al. 1996). However, the thermally softened and attenuated Cuyania slab, considered by itself, seems unlikely to have been the colliding body owing to its limited size, diminished coherence, and increased buoyancy relative to the crust of the Gondwanan continental mass (Cloos 1993; see also Buggisch 1996). In addition, the Ordovician contractional tectonism in the Western Sierras Pampeanas occurred during the Ocloyic event, significantly later than Late Cambrian to early Caradoc deformation and magmatism in the Precordillera/Marathon/Solitario basins.

Conclusions The Precordillera, an element of the Cuyania terrane of western Argentina, originated in the Ouachita embayment of the southern Laurentian margin. The Precordillera and the Marathon/ Solitario basin of west Texas, both with fundaments of Grenvillian Laurentian basement, evolved together during Late Cambrian through to mid-Ordovician times. The Marathon/Solitario basin was a twosided basin, not merely an embayment in the

THE PRECORDILLERAN ODYSSEY passive margin. To the north it was bounded by the circum-Laurentian carbonate platform, from which it received much carbonate detritus. Siliciclastic/volcaniclastic sands within the early Palaeozoic succession came from a southerly source, believed to have been recta-igneous basement rocks in Cuyania, with probable contributions from the Famatina eruptive complex. Both Early Ordovician carbonate platforms were home to virtually identical reef organisms. The sponge-algal facies that hosts the Pulchrilamina (west Texas) and Zondarella (Precordillera) reefs was mainly distributed around the Ouachita margin of Laurentia. Syn-sedimentary extensional block-faulting characterized the conjoined basin throughout its existence. Arenig-Llanvirn meta-bentonites and ash-rich pyroclastic flows of the Marathon/ Solitario/Precordillera basin were probable products of that extensional deformation. LlandeiloCaradoc igneous activity reflects the culmination of extension and extrusion of pillow basalts in turbidites of the westernmost Precordillera. Significant unconformities in both the Marathon/ Solitario and Precordilleran sections span the period from later early Caradoc until early Ashgill - the time of final separation. Marathon/Solitario/Precordillera extension, sedimentation, and volcanism are interpreted as having occurred within a right-transtensional regime created by northward translation and clockwise movement of Laurentia with respect to Gondwana. With continued extension and right-oblique separation of the two continental masses, the attenuated and thermally softened Laurentian slab broke apart and oceanic tholeiites were emplaced in turbidites of the western Precordillera. Severance of Cuyania from Laurentia was thus complete before the onset of Taconic or Ocloyic orogenesis. This manuscript has benefited from thoughtful reviews by W, R. Muehlberger, C. W. Rapela, and particularly that ofW. A. Thomas. Lively discussions in the field in the Solitario with V. A. Ramos, L. H. Dalla Salda, R. Astini, I. W. D. Dalziel, W. R. Muehlberger. F. Hutson, and S. M. Kay helped shape the interpretations presented here. Graptolites collected by Ramos, Dalziel and Dickerson on that occasion were identified by W. B. N. Berry, whose collaboration is most gratefully acknowledged. Dickerson thanks S. Tweedy (formerly of Mineral Studies Laboratory, Texas Bureau of Economic Geology), B. Hanson (Coming Laboratories), R. Folk, L. Lynch, D. S. Barker, K. Milliken, F. Hutson (all of University of Texas at Austin) for edification on clay mineralogy/ diagenesis, interpretation of XRD data, melt inclusions, igneous geochemistry, and mineral separation techniques. Fieldwork with W. Muehlberger in the

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Solitario has been endlessly enlightening. Stimulating exchanges with J. D. Cooper, G. W. Viele, H. Bahlburg, D. L. Amsbury, V. A. Ramos, I. W. D. Dalziel, and O. Lehnert in various venues have expanded Dickerson's views of the Precordillera, west Texas, the Ordovician, Argentina and the world in general.

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Ordovician K-bentonites in the Argentine Precordillera: relations to Gondwana margin evolution WARREN

D. H U F F ~, S T I G M. B E R G S T R O M 2, D E N N I S

CARLOS

A. C I N G O L A N I 4 A N D

RICARDO

R. K O L A T A 3,

A. A S T I N I 5

1Department o f Geology, University of Cincinnati, Cincinnati, O H 45221, USA 2 Department o f Geological Sciences, The Ohio State University, 155 S. Oval Mall, Columbus, O H 43210, USA 3 Illinois State Geological Survey, 615 E. Peabody Dr., Champaign, IL 61820, USA 4 Centro de Investigaciones Geol6gicas, Universidad Nacional de La Plata, Calle 1, No. 644, 1900 La Plata, Argentina 5 CONICET, C(ttedra de Estratigrafia y Geologia Hist6rica, Facultad de Ciencias Exactas, Fisicas y Naturales, Universidad Nacional de C6rdoba, Av. V~lez S~rsfield 299, CC#395, 5000 C6rdoba, Argentina Abstract: Ordovician K-bentonites have now been recorded from >20 localities in the vicin-

ity of the Argentine Precordillera. Most occur in the eastern thrust belts, in the San Juan Limestone and the overlying the Gualcamayo Formation, but a few ash beds are known also from the central thrust belts. The oldest occur in the middle Arenig I. victoriaelunatus graptolite (Oe. evae conodont) Zone, and the youngest in the middle Llanvirn P. elegans (P. suecicus) Zone. Mineralogical characteristics, typical of other Ordovician K-bentonites, include a matrix of illite/smectite mixed-layer clay and a typical felsic volcanic phenocryst assemblage: biotite, beta-form quartz, alkali and plagioclase feldspar, apatite, and zircon, with lesser amounts of hornblende, clinopyroxene, titanite and Fe-Ti oxides. The proportions of the mineral phases and variations in their crystal chemistry are commonly unique to individual (or small groups of) K-bentonite beds. Glass melt inclusions preserved in quartz are rhyolitic in composition. The sequence is unique in its abundance of K-bentonite beds, but a close association between the Precordillera and other Ordovician sedimentary basins cannot be established. The ash distribution is most consistent with palaeogeographical reconstructions in which early Ordovician drifting of the Precordillera occurred in proximity to one or more volcanic arcs, and with eventual collision along the Andean margin of Gondwana during the mid-Ordovician Ocloyic event of the Famatinian orogeny. The Puna-Famatina terrane northeast of the Precordillera might have served as the source of the K-bentonite ashes, possibly in concert with active arc magmatism on the Gondwana plate itself.

Lower Palaeozoic K-bentonites (altered volcanic ashes) are widely distributed on the continents bordering the former Iapetus Ocean, especially in northwestern Europe and eastern North America. Local and regional studies during the last decade have added a wealth of new information about their geographical and stratigraphical distribution, geochemistry, and mineralogy as well as their tectono-magmatic and palaeogeographical significance (Kolata et al. 1996; Bergstr6m et al. 1997; Huff et al. 1998). In South America, K-bentonites are as yet unknown in the Cambrian and Silurian and were discovered in the Ordovician only a few years ago (Huff et al. 1995a). During the past few years, many Lower and Middle Ordovician K-bentonite beds have been found in the Precordillera of western Argentina; indeed, this region has some of the most abundant ash beds

known from that period anywhere in the world. K-bentonite geochemistry may provide highly significant information about the tectono-magmatic nature of the source area, and the distribution patterns of individual ash beds or complexes of such beds may shed additional light on the former positions of continental plates. Specific chemical and mineralogical features of primary phenocrysts can serve as important criteria for regional stratigraphical correlation of beds and bed sequences, whether expressed as elemental ratios of individual grains or in the immobile element chemistry of bulk K-bentonite samples (Kolata et al. 1987; Huff et al. 1992; Bergstr6m et al. 1995). K-bentonite beds also provide important evidence regarding the timing and location of orogenic events because they typically originate from source volcanoes situated at or near tectonically active plate margins (Kolata et al.

HUFF, W. D., BERGSTR6M, S. M., KOLATA, D. R. et al. 1998. Ordovician K-bentonites in the Argentine Precordillera: relations to Gondwana margin evolution. In: PANKHURST,R. J. & RAPELA,C. W. (eds) The ProtoAndean Margin of Gondwana. Geological Society, London, Special Publications, 142, 107-126.

108

W. D. HUFF E T A L .

1987; Huff et al. 1992; Bergstr6m et al. 1995). Further, dating of K-bentonite beds or complexes of beds provides precise age dates on periods of volcanism associated with continental margin subduction events. Therefore, when trying to unravel the pre-Andean evolution of the western Gondwana margin it is appropriate to consider the various types of evidence provided by the K-bentonites. The first step in any such evaluation is to explore the vertical and horizontal distribution of these ashes, both locally and regionally, and to clarify their age and distribution patterns. Without establishing the position of these beds in a reliable biostratigraphical framework, no meaningful regional comparisons can be made, and the assessment of their palaeogeographical significance is seriously hampered. In view of this, the first part of the present contribution will assess the stratigraphical and geographical distribution of the Ordovician K-bentonites in the Argentine Precordillera. We will also make a comparison with the geographical and stratigraphical distribution of Ordovician K-bentonite complexes in North America (Laurentia) to examine whether or not they are likely to represent the same ash falls. If they do, it would obviously suggest rather close proximity of Laurentia to the Precordillera during the time of the ash fall (Dalziel 1997). The absence of correlative beds on these two land masses would be consistent with models which argue for a wide Iapetus at the time of eruption (Astini et al. 1995; Thomas & Astini 1996), unless prevailing wind direction severely skewed ash distribution. Thus, K-bentonites may serve as a means of testing palaeogeographical reconstructions and models of the Iapetus Ocean. A second part of the present contribution deals with the mineralogy and geochemistry of Precordilleran K-bentonites and their tectonomagmatic significance. We show that melt inclusion composition and the whole rock ratios of immobile elements provide evidence for collision-margin volcanism. A final part summarises the evidence at hand bearing on the evolution of the Gondwana margin, and the Ordovician geological history of the Precordillera. Our investigations of the Argentine Ordovician K-bentonites are continuing and some of the data presented below are of a somewhat preliminary nature.

Occurrence Since their discovery in the Precordillera in 1994, Ordovician K-bentonites have been recorded from more than 20 localities in a region extending about 250 km in a north-south direction across

parts of San Juan and La Rioja Provinces. Some of the principal localities are shown in Fig. l. No ash beds have been observed in the Ordovician outcrops in the San Rafael region some 200 km south of Mendoza. Likewise, as noted by Bergstr6m et al. (1996), no K-bentonite beds are known from the locally rather strongly tectonized slope and basin successions in the western Precordillera, such as the well-exposed Middle Ordovician sequences near the northern end of Sierra de la Invernada about 40km northeast of Talacasto (Fig. 1), or in the outcrops along the mountain front from Calingasta northwards to the Jfichal River. Most of the known K-bentonites occur in the eastern thrusts of the thrust-and-fold belt, where they are quite common in the upper section of the San Juan Limestone, characterized by massive platform carbonates with an early Palaeozoic Laurentian fauna, and in the overlying deep water graptolitic shales of the Gualcamayo Formation. A few ash beds are also known from the central thrusts. For convenience, we will briefly review the geographical and stratigraphical distribution patterns in terms of three main distribution areas, namely the Guandacol region in the northernmost Precordillera, the Jfichal region some 100 km to the south, and the San Juan region about 150kin still farther to the south (Fig. 1). Only a few key sections are discussed here; for further information, see Bergstr6m et al. (1996). The oldest recorded Palaeozoic K-bentonites in the Precordillera occur in the middle Arenig Oe. evae Zone (Htinicken & Sarmiento 1982, 1985) in the topmost part of the Lower Ordovician San Juan Limestone in outcrops along the Gualcamayo River in the Guandacol region. In the overlying, about 200m thick, Gualcamayo Formation, which ranges from the Upper Arenig to at least the Middle and probably Upper Llanvirn, there are numerous ash beds, especially along a small tributary to the Guandacol River (26~ 68~ where more than 170 separate ash beds have been observed. This is by far the most extensive suite of Ordovician K-bentonites recorded from an outcrop anywhere in the world and ranks slightly above the Lower Silurian K-bentonite succession at Dob's Linn in Scotland which contains 135 beds (Merriman & Roberts 1990). Most (c. 85) of the ash beds, which range in thickness from less than 1 mm to more than 50 cm, occur in the upper Arenig U. a u s t r o d e n tatus Zone (Brussa & Astini 1997) but many beds are present also in slightly older and slightly younger strata. Excellent sections of coeval strata, but with fewer K-bentonite beds, are present at the Quebrada de Los Saltitos and

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in the Quebrada Nazareno, a short tributary branch between Quebrada Las Plantas and Quebrada de Potrerillos in the Guandacol area (Fig. 1). In the Jfichal area, there are about 30 individual K-bentonite beds in the Los Azules Formation (equivalent to the Gualcamayo Formation elsewhere) at Cerro Viejo (Huff et al. 1995a; Bergstr6m et al. 1996; Cingolani et al. 1997) and up to a dozen such beds have also been recorded from the upper part of the underlying San Juan Limestone. A concordant U - P b zircon age from Cerro Viejo was reported by Huff et al. (1997b) as 464 i 2 Ma which they considered to be the age of the base of the U. austrodentatus Zone. Beds of the latter complex, which post-date the oldest K-bentonite beds in the Guandacol region, are well exposed at many localities, such as Cerro Potrerillo, Cerro La Chilca, and La Silla (Fig. 1), and also in the still poorly dated successions at Las Chacritas (30~ 68~ (Albanesi & Astini 1994; Astini 1994) and Mogotes Azules (30~ 68~ Multiple K-bentonite sequences are present at all of these localities but none are as extensive as the middle Arenig K-bentonite complex in the Guandacol region. The most accessible K-bentonite locality in the San Juan region is in a highway road cut at Quebrada de Talacasto about 55km northnorthwest of San Juan. As described by Bergstr6m et al. (1996) there are 30 individual ash beds, the thickest ones 10-15 cm, within a 1101n thick succession of the upper San Juan Limestone. At least the upper portion of this complex is probably comparable with the K-bentonite complex in the Gualcamayo Formation at Cerr~ Viejo. The considerable thickness of some beds at the latter locality may be due to the fact that the section is more condensed and some of the thick beds may be composed of several, in time relatively closely spaced, ash falls that are represented by separate ash beds in the thick Talacasto succession. A single bed about 50m below the top of the San Juan Limestone at the Tambolar section in the San Juan River valley (Lehnert 1995; Bergstr6m et al. 1996) is one cf the oldest K-bentonite beds known from the Precordillera. Fig. 1. Location map of the Precordillera, Argentina, showing some of the principal localities where K-bentonites are exposed. The inset map of Argentina shows the position of the Precordillera. The areas in black are the Ordovician outcrop areas of the eastern and central thrust belts, and the stippled areas represent the Ordovician outcrop areas of the western thrust belts. Numbered locations refer to the principal sections used in this study.

110

W. D. HUFF E T A L . or Upper Ordovician strata9 Figure 2 is a n o r t h south transect showing the main patterns of K-bentonite bed distribution. The high concentration of ash beds in the upper Arenig-lower Llanvirn in the Guandacol region may suggest the general direction toward the location of the

In terms of stratigraphical distribution, the oldest K-bentonites currently known in the Precordillera are of middle Arenig age and the youngest are of middle to upper Llanvirn age (Bergstr6m et al. 1996)9 No ash beds have been observed in younger parts of Middle Ordovician

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PRECORDILLERA K-BENTONITES source area; alternatively, it may reflect the fact that these ash beds are in deeper-water dark shales that represented an environment which was more favourable for preservation of volcanic ashes than the shallower-water carbonate dominated depositional environments of the San Juan Limestone in the Jfichal-San Juan regions. In terms of east-west geographical distribution, there is an obvious trend from many beds in the eastern thrust belts to relatively few in the central thrust belts (Fig. 3). In many sections in the central belts, the upper San Juan Limestone and the overlying Gualcamayo Formation are not preserved, being cut out by an unconformity (Astini et al. 1995; Lehnert 1995b). Hence, in those sections the K-bentonite rich interval of the eastern belt may not be represented. Elsewhere, Silurian strata overlie younger Ordovician units, even in the central ranges (Astini & Maretto 1996)

Ordovician volcanism The widespread occurrence of K-bentonite beds in the Argentine Precordillera constitutes one of the most extensive suites of such beds known anywhere in the Ordovician System of the world and serves as testimony to the high intensity of explosive volcanism along this margin of Gondwana during the early and middle parts of that period. Previous and ongoing studies of the sedimentology, mineralogy and geochemistry of these beds provide both insight and constraints concerning the magmatic, tectonic and paleogeographical settings under which the explosive volcanism was generated, and also permit comparisons with lower Palaeozoic K-bentonites on other continents. While our most recent field work has revealed an extensive succession of K-bentonite beds in the exposures along the Gualcamayo River and its tributaries in western La Rioja Province, most of our detailed particlesize, mineralogical, and geochemical studies to date have been on samples from the extensive sections at Cerro Viejo, near Jfichal, and at Talacasto, north of San Juan, in San Juan Province (Fig. 1). Further, while most known evidence for pre-Andean explosive volcanism on the Gondwana margin is preserved in the Ordovician sections of the Argentine Precordillera, additional beds of pyroclastic origin have also been reported from the Balcarce Formation of the Tandilia region, south of Buenos Aires (Dristas & Frisicale 1987). This formation is considered to be Ordovician in age based on Cruziana ichnofaunal evidence and a K - A r radiometric age of 396 + 11 Ma from a cross-cutting dyke (Cingolani

111

et al. 1985). At least one, and perhaps as many as

four, altered pyroclastic beds occur in the white quartzite sequence which ranges from 18 to 500 m in thickness and unconformably overlies Precambrian basement (Dalla Salda et al. 1988). In contrast to the illite/smectite-rich beds of the Precordillera, the Balcarce beds consist mainly of well crystallized kaolinite, indicated by sharp X-ray powder diffraction reflections at 7.12A and 3.58A, with occasional crystals of altered ilmenite, and are considered to be the product of altered mafic ashes (Dristas & Frisicale 1987). They may legitimately be called tonsteins because of their clay mineralogy and apparent pyroclastic origin, although, in contrast to most tonsteins, they occur in quartzites and have no association with coal measures. Their relationship to the Precordillera K-bentonites, if any, is unclear, other than that they also appear to represent evidence of fallout ash accumulation sometime during the Ordovician. The Balcarce beds most probably represent an entirely different tectonic setting from that of the felsic Precordillera beds, which will be the focus of the remaining discussion.

K-bentonite mineralogy Precordillera K-bentonites consist principally of authigenic clay minerals, carbonates, sulphates, oxides, and sulphides, plus primary magmatic quartz, titanite, apatite, biotite, feldspar, pyroxene, amphibole, zircon and Fe-Ti oxides (Huff et al. 1995a; Krekeler et al. 1995; Cingolani et al. 1997; Prokopenko et al. 1997). The clays and other secondary minerals provide information about the alteration processes that have affected the fallout ash, and the primary pyroclastic crystals (herein called phenocrysts) retain information concerning composition and origin of the parental magma as well as the timing and tectonic setting of the source volcanism. The primary mineral suite is characteristic of the felsic calc-alkaline nature of the tephras from which these, as well as the majority of Lower Palaeozoic K-bentonites, are derived. Clay mineralogy

The clay mineral fraction of Precordilleran Kbentonites consists of R1 to R3 ordered mixedlayer illite/smectite (I/S) with the illite content ranging from 72 to 95% (Fig. 4). The X-ray diffraction (XRD) traces shown in Fig. 4 are representative of more than 50 such analyses of oriented, glycol-saturated, < 2 # m size fractions of clay samples from Talacasto, Cerro Viejo,

112

W . D . H U F F ET AL. been smectite. Kaolinite is also present in some K-bentonites, at times in nearly equal proportion to the I/S. However, kaolinite is less common in Ordovician K-bentonites than in Silurian K-bentonites, in which it is common and, even occasionally, the dominant clay mineral (Teale & Spears 1986; Huff et al. 1997c). In the Precordillera, Lehnert & Keller (1994) have recorded very low-grade metamorphic temperatures based on conodont alteration index (CAI) studies in the central and western zones. Most K-bentonites in the eastern thrust belts, however, were not subjected to temperatures greater than about 100~'C (Lehnert & Keller 1994; Ortega et al. 1996).

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~174 Fig. 4. Powder X-ray diffraction patterns of oriented. glycol-saturated slides of the <2 #m fraction of eight representative Precordillera K-bentonites. All contain illite/smectite as the principal clay mineral and are arranged in sequence from R1 ordered I/S with 72% I at the base to R3 ordered I/S with 95% I at the top as indicated by peaksobetween 9.9-12.9 A. Low-angle reflections at 33.3 A indicate RI ordering. Percentages were calculated using NEWMOD (Reynolds 1985). Samples 5B & 6 are from Cerro Viejo, 32 is from La Silla, 47, 50 & 52 are from Talacasto, and 22 & 23 are from Quebrada Las Plantas at Guandacol. Kaolinite is a major constituent in several samples and can be identified by peaks at 7.14,~ and 3.58A. Guandacol, and La Silla. The relative proportions of illite and smectite as well as the nature of their ordered interstratification were determined using the computer program NEWMOD~ (Reynolds 1985). The traces in Fig. 4 are stacked to show the range of illite found. With few exceptions, I/S is the dominant clay mineral component of Lower Palaeozoic K-bentonites throughout the world, and reflects both the vitric-rich nature of the original tephra and also their calc-alkaline character. The relative proportions of illite and smectite layers are a function of several factors, including thermal history, extent of K-metasomatism, availability of potassium and aluminium in the system, and initial composition of the glass alteration products, generally considered to have

Figure 5 shows examples of primary volcanogenic phenocrysts from bed ARG-3 at Cerro Viejo, which are typical of many Precordilleran K-bentonites. Quartz is very common and frequently occurs in beta-form morphology, characteristic of volcanogenic quartz (Bohor & Triplehorn 1993). Individual crystals average 0.2 mm in diameter, based on optical measurements of over 400 grains from four separate K-bentonite beds at Cerro Viejo (Fig. 6). Previous studies (Huff et al. 1996) of K-bentonite phenocrysts report that despite the loss of original pumice and glass particles during diagenesis, some phenocrysts remain unaffected and can serve as reliable proxies for original grain size. Thus~ it may be possible to speculate on some dynamic aspects of ash emplacement, such as the eruptive mechanism, by comparing grain-size measurements with published theoretical models of eruptive plume dispersion. Quartz was selected for size measurement because of its relatively equidimensional character and thus the greater likelihood that its average diameter would more closely reflect transport conditions for ideal spherical grains (Walker 1981). The exponential decrease in grain size of fallout tephra as a function of distance from the source is well known (Carey et al. 1989; Sigurdsson & Carey 1989) and holds equally well for co-ignimbrite and plinian plumes. A comparison of the Cerro Viejo grain-size data with theoretical models suggests that particles averaging 0.2 mm in diameter were transported at least 200 km from the source vent or caldera. This assumes, for example, a minimum plume height of 35 km and a sustained wind velocity of 50 km/hr, both of which are well within the range of recent large-scale explosive eruptions. The absence of coeval volcanism anywhere on the Precordillera further suggests the source volcanoes were not in the immediate

PRECORDILLERA K-BENTONITES

113

Fig. 5. Examples of volcanogenic phenocrysts from bed ARG-3 at Cerro Viejo. (a) SEM photograph of a zircon crystal. The scale bar is 100 Hm. (b) Beta-form quartz crystals averaging 0.2 mm in diameter. Note the clear, round glass melt inclusion in the centre crystal. (e) Beta-form quartz crystal with a sphene inclusion. (d) A zircon (shorter) and apatite (longer) crystal, both containing elongated melt inclusions. Note internal zoning in the zircon.

vicinity. A likely source is the Puna-Famatina terrane to the north (Bahlburg & Herv~ 1997) which appears to have had a geodynamic history similar in age and style to the Precordillera. Some quartz crystals contain pristine glass melt inclusions, which are clear, rounded or ovoid in shape, and range between 20 and 75 #m in diameter. We report below the results of microprobe analyses of some melt inclusions which serve to constrain the composition of the pre-eruption parental magma. Some grains also contain inclusions of titanite and opaque minerals. Zircon and apatite are both common constituents of Precordillera K-bentonites. Zircon crystals are euhedral and typically 100 200 # m in length, with high aspect ratios. Some show internal zoning and also contain elongated melt inclusions (Fig. 5). Apatite grains tend to be somewhat larger than zircons and also contain melt inclusions. Pyroxene was identified in at least five Kbentonite beds at Talacasto where they constitute as much as 80% of the non-opaque heavy mineral fraction (Prokopenko et al. 1997). The crystals appear as angular, highly fragmented grains ranging from pale green to greenish-brown

in colour under the petrographical microscope. Microprobe analyses (Fig. 7) show that they are calcium-rich clinopyroxenes with a diopside composition. Their occurrence is somewhat unusual when compared to the North American and Baltoscandic K-bentonites which do not have them, but it is consistent with the general mineralogy of volcanic rocks generated in silicic magma chambers. For example, Hildreth (1979) reported both orthopyroxenes and clinopyroxenes from the Bishop Tuff, a layered silicic ashflow tuff. The directions of compositional trends ofclinopyroxenes from the Talacasto section present some systematic changes that appear to vary with stratigraphical position. For example, augites with the least Ca-enrichment from the two lowest beds in the section (Fig. 7) display a tholeiitic differentiation trend (Sack & Ghiurso 1994) whereas crystals from the uppermost parts of the section are much more clustered and show little compositional variation, or show trends along the Ca-Fe line of the clinopyroxene quadrilateral. Clinopyroxenes crystallizing from felsic magmas would be expected to show enrichment in Fe and Ca if mixing with more mafic magmas

114

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were involved. Other variables, such as changes in oxygen fugacity, which will affect the amount of Fe in the clinopyroxene structure, and assimilation of continental crustal rocks, which might be expected to dilute the Ca-Fe content, might also influence the observed variations. Beds in the upper part of the section contain both high-Ca and low-Ca clinopyroxenes which may reflect coexisting crystallization of both phases. Thus, pyroxene compositions in the Talacasto section suggest a somewhat more complex history of magmatic development than previously shown by other Ordovician K-bentonites in the Precordillera or elsewhere along the Iapetus margin. Such information should ultimately prove quite valuable in interpreting the tectonic history of this region of the Gondwana margin. Feldspars constitute one of the more common phenocryst phases in Precordilleran K-bentonites, in part because they occur as both primary and authigenic minerals. Feldspars normally make up to as much as 20% of the total grain

population and range from 0.2mm, for the majority of grains, occasionally up to 0.5mm. The larger grains appear to be mainly authigenic albite, however, while the smaller grains are characteristically primary plagioclase and K-feldspar. Preliminary microprobe studies of the feldspar population of several beds at Cerro Viejo (Krekeler et al. 1995) revealed a broad range of compositions. In particular, two K-bentonites, ARG9 and ARG10, have anomalously high percentages of volcanogenic feldspar, constituting nearly 40% of the total grain population. Microprobe analyses indicate that the plagioclase compositions differ from bed to bed and range in composition from An31 to Ans0. Sanidine compositions were obtained from a single K-bentonite, ARG12, where they range from Or26 to Or19. A ternary diagram of all feldspar data, including authigenic grains, is presented in Fig. 8. This range of feldspar compositions provides additional evidence of the heterogeneous nature of the source magmas.

PRECORDILLERA K-BENTONITES

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Fe-oxides occur commonly in the K-bentonites and have been identified using reflected light microscopy and X-ray powder diffraction. Petrography reveals that some Fe-oxides have nucleated on other opaque phases such as ilmenite and magnetite and were most probably derived from them by alteration. The Fe-oxides occur in a variety of crystal habits including tetragonal

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bipyramids, isometric cubes, and rounded radiating grains. Many rounded grains have crystal faces developed on the surface: X-ray diffraction shows them to be dominantly goethite.

Geochemistry Magmatic and tectonic discrimination diagrams can be used to infer original magma composition and source volcano settings relative to plate margins. Microprobe analysis of glass melt inclusions in quartz grains from Cerro Viejo (Table l) yielded oxide totals near 96.6%, and the data were plotted on a total alkalis versus silica (TAS) diagram (Fig. 9). Care was taken to minimize Naqoss by reducing the beam current, enlarging the spot size, and averaging repeated analyses (Spray & Rae 1995; Hanson et al. 1996) so it is considered that most of the remaining 3.4% consists of H20. This conclusion is supported by previous studies of glass melt inclusions in which total O was measured and the major deficit in the oxide total was found to be H20 (Nash 1992). Silicic magmas may also contain some exsolved volatile constituents which would have been released at the time of eruption and thus would not be preserved in the trapped glass (Wallace et al. 1995). Felsic magmas frequently contain dissolved water in the range of 3-5 wt%, and are capable of producing large-scale explosive eruptions, particularly when associated with caldera-forming events (Rose & Chesner 1987; Gardner et al. 1991).

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Whole-rock major and trace element analyses of selected Precordillera K-bentonites are given in Table 2. Chondrite-normalized rare earth element (REE) plots of five Cerro Viejo samples (Fig. 10) are relatively enriched in light REE and relatively depleted in heavy REE, which is typical of highly evolved, felsic magmas. Moreover, most have a

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Y (ppm) Fig. 11. A tectonic discrimination diagram from Pearce et al. (1984) showing the position of the Cerro Viejo volcanic rocks in terms of granitic origins. WPG, within plate granite; ORG, ocean ridge granite; VAG, volcanic arc granite; syn-COLG, syn-collision granite. The samples fall on the boundary between volcanic arc and within plate granites, typical of collision margin felsic volcanic rocks. fairly pronounced negative Eu anomaly, characteristic of felsic magmas in which most of the plagioclase has already crystallized. When plotted on the granite tectonic discrimination diagram of Pearce et al. (1984), the Cerro Viejo volcanic rocks fall on the boundary between volcanic arc and within plate granites (Fig. 11). Based on comparison with previously-studied K-bentonite beds (Bergstr6m et al. 1995; Huff et al. 1996; Kolata et al. 1996), these appear to represent vitric fallout ash from large-scale explosive volcanism associated with collision-margin tectonism. These data are consistent with the presence of an active Gondwana continental margin in the Early-Mid-Ordovician (Coira et al. 1982; Ramos et al. 1986; Dalla Salda et al. 1992a).

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An unresolved question concerning the Precordilleran K-bentonites is the location of the volcanoes that produced these ashes. In the case of Quaternary and Holocene ashes, regional mapping combined with geochemical and mineralogical fingerprinting is frequently capable of pinpointing the location of the source vents. However, with ancient volcanic rocks those source vents have generally been destroyed long since and one is confronted with problems similar to most palaeogeographical reconstructions.

118

W. D. H U F F E T AL.

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Assessment of variation in the number of beds, their stratigraphical distribution, and their relative thickness within a region, may provide some clues to the direction of ash transport (Huffet al. 1992, 1996; Kolata et al. 1996). Such assessment in the Precordillera, however, is complicated by the fact that the K-bentonites occur in thrust belts that were subjected to eastward transport of large but uncertain distance during Andean tectonism. A 50% crustal shortening was estimated along the San Juan River by von Gosen (1992), whereas Allmendinger et al. (1982) calculated nearly 95% shortening along the Jfichal River. Moreover, recent work has shown that the top of the San Juan Limestone is of different age at different localities (Htinicken & Sarmiento 1985; Astini et al. 1995; Lehnert 1995a) Locally, it is marked by a very prominent, largely pre-Silurian, unconformity; hence, the interval in the uppermost San Juan Limestone and the overlying Gualcamayo (or Los Azules) Formation that contains the most numerous K-bentonites may be very poorly, if at all, represented in some sections.

A regional comparison of global Ordovician K-bentonite complexes is shown in Fig. 12 and provides some constraints on the possible juxtaposition of various tectonic plates and plate segments at the time of tephra formation. For example, the abundance and stratigraphical similarity of Caradoc beds in North America and Baltoscandic prompted Huff et al. (1992) to propose that these two regions were in close proximity at that time. Haynes et al. (1995) challanged the proposed correlation of the Millbrig and Kinnekuile beds on the basis of an apparent dissimilarity in biotite compositions. However, current studies (Huff et al. 1997a) suggest that the Millbrig beds consist of multiple units, at least one of which is chemically indistinguishable from the Kinnekulle bed. The Argentine sequence, on the other hand, has a different age distribution of K-bentonite beds, and they provide no supporting evidence of a close association between the Precordillera and the other Ordovician sedimentary basins in Fig. 12.

PRECORDILLERA K-BENTONITES Comparison

with North America

More than 60 K-bentonite beds are currently known from the Ordovician of North America (Kolata et al. 1996) and some of these beds, such as the Deicke and the Millbrig K-bentonites, are widely distributed. However, a comparison of the stratigraphical distributions of ash beds in the Precordillera and North America (Fig. 13) show

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considerable differences (Bergstr6m et al. 1996). In the latter region, very few ash beds are known from Lower Ordovician and the lower MidOrdovician. These include single beds in Mississippi (Thomas 1988; Kolata et al. 1996) and in the Marathon area, western Texas (King 1937), an undetermined number of beds of early MidOrdovician age in a giant sinkhole near Douglas Lake, eastern Tennessee (Laurence 1944), and a few ash beds in western Newfoundland (Kolata et al. 1996). Keller & Dickerson (1996) recently misidentified some calcareous mudstones of early Mid-Ordovician age from the Solitario region, western Texas, as K-bentonites. Hence, we currently know of no obvious counterpart in North America to the extensive K-bentonite bed complex of Early and early Mid-Ordovician age in the Precordillera. However, it may be significant to note that much of the lower Middle Ordovician is not represented in the successions in the Southern and Central Appalachians, where this interval is cut out by the prominent Sauk-Tippecanoe unconformity (Ross et al. 1982; Faill 1997). The principal K-bentonite complex in eastern North America is in the upper Middle Ordovician (Mohawkian) where there are more than 40 ash beds (Bergstr6m et al. 1996; Kolata et al. 1996), two of which may reach a thickness of 0.1-1 m (Huffet al. 1996). No equivalents to this K-bentonite complex, nor to the few North American Upper Ordovician ash beds, are known in the Precordillera. This can be taken as an indication that the close Mid-Ordovician juxtaposition of the Precordillera and easternmost North America shown in several recent paleogeographical reconstructions (Dalla Salda et al. 1992a; Dalziel et al. 1994) is unlikely to be correct. On the other hand, juxtaposition of North America and Baltoscandia is more likely for this time interval, as suggested by Huffet al. (1995b).

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magmatic

arc

Based on the geochemical characteristics of Precordilleran K-bentonites, the volcanoes from which the ashes erupted developed on continental crust in a convergent tectonic setting, either in the form of an arc situated on the Gondwana margin or as a colliding terrane of continental affinity. Several authors interpret the Eastern P u n a - F a m a t i n a magmatic belt as an outer magmatic arc (Rapela et al. 1992) that was located close to the southwestern Gondwana margin in the south Iapetus Ocean (Dalziel et al. 1994). Palaeomagnetic evidence lends support to this hypothesis by placing the Eastern P u n a - F a m a tina belt in the south Iapetus Ocean, close to

122

W. D. HUFF E T AL.

the Gondwana margin, during Mid-Ordovician time (Conti et al. 1996). Palaeobiogeographical arguments (Benedetto & Sfinchez 1996) further support an intra-Iapetus positioning for the Famatina Ordovician series. The accretion of this arc to the Gondwana craton was accompanied by the emplacement of a number of granitoid plutons in the Western Sierras Pampeanas. These plutons were loosely identified as pre-Taconic, Taconic, and post-Taconic in age (Lopez de Luchi & Dalla Salda 1995) in reference to the geodynamic model proposed by Dalla Salda et al. (1992b), which envisaged the Taconic Appalachians and the Famatinian arc as a single orogenic belt. The same plutonic groups had been labelled G 1, G2, and G3, respectively, by Rapela et al. (1990). The pre-Taconic or G1 group (Late Cambrian-Early Ordovician granitoids, 520460 Ma) are stocks connected to minor granitic plutons. The Taconic or G2 granitoid group (460-420Ma) includes plutons concordant to regional foliation, essentially biotite-muscovite peraluminous monzogranites to syenogranites, with minor granodiorites and tonalites associated with an extensive granitic migmatization. The post-Taconic or G3 granitoids (420-360 Ma) are peraluminous to slightly peralkaline and poorer in MgO than the Taconian granitoids. The onset of granitoid emplacement and associated volcanism coincided with the development of sedimentary basins in the northern and central Sierra de Famatina and in the Puna Palaeozoic belt (Coira et al. 1982; Koukharsky et al. 1988; Coira & Nullo 1989). A more recent evaluation of these ages is provided by Pankhurst et al. (this volume). In the Puna (Altiplano) a pyroclastic-sedimentary marine sequence of Arenig to Llanvirn age can be recognized, and geochemical data from these volcaniclastic rocks suggest a continental arc source. Mannheim (1993) concluded that the geochemical features of the early Palaeozoic rocks in the Sierra de Famatina resulted from two different environments of magma generation. In one case, subduction-related basic primary magmas were generated in an island arc setting, and in the second case, Tremadoc to Llanvirn acidic peraluminous magmas formed during crustal anatexis and produced pyroclastic deposits such as Las Planchadas Formation. Toselli et al. (1990) described an episode of Ordovician magmatism and volcanism in which pyroclastic materials (>400km 2) were derived from acidic to intermediate magmas. This magmatism was synchronous with sedimentation of shallow water rocks of early-mid-Arenig age (Acefiolaza & Toselli 1984; Vaccari et al. 1993). Although these Famatinian rocks overlap in age with strata containing K-bentonites in the Guandacol

region, they appear to contain no K-bentonites or similar pyroclastic deposits. However, further comparative studies, such as geochemical fingerprinting, precise age dating, and mineralogical investigations, are still needed to clarify many aspects of the tectono-thermal setting and volcanic source of Lower-Middle Ordovician Precordilleran K-bentonites.

Discussion and conclusions Numerous K-bentonites occur in the San Juan and Gualcamayo Formations in the eastern thrust belts of the Precordillera, and a few ash beds are also known from the central thrust belts. They range in age from the middle Arenig L victoriae lunatus graptolite (Oe. evae conodont) Zone to middle Llanvirn age in the P. elegans (P. suecicus) Zone. As is typical of many Ordovician K-bentonites found elsewhere, they consist of a matrix of illite/smectite mixedlayer clay plus a phenocryst assemblage including biotite, beta-form quartz, alkali and plagioclase feldspars, apatite, and zircon, hornblende, clinopyroxene, titanite, and Fe-Ti oxides. Pristine melt inclusions in quartz crystals are rhyolitic in composition. The Argentine sequence is unique in its abundance of K-bentonite beds, yet when compared with K-bentonite sequences of similar age elsewhere it provides no supporting evidence of a close association between the Precordillera and other Ordovician sedimentary basins along the Iapetus margin at that time. Palaeogeographical reconstruction for the early Palaeozoic is, at best, a 'weight of evidence' process, in which the more pieces of information there are available, the more solid the arguments for a particular scenario. K-bentonite stratigraphical and geographical distribution patterns can be helpful in this process by providing direct evidence of subduction-related explosive volcanism and the accompanying palaeo-wind directions, as well as the relevant sedimentary conditions and magmatic processes. Based on that record, it is clear that subduction-related volcanism in the Precordillera commenced in the mid-Arenig and continued to the late-Llanvirn, coeval with volcanism in the Sierra de Famatina. It has been proposed that the Ocloyic event of the Famatinian orogeny in South America was a continuation of the Taconic orogeny in North America, based on similarities between Cambrian to early Ordovician carbonate sequences in the Precordillera and those of sections along the Laurentian margin (Dalla Salda et al. 1992b; Dalziel et al. 1994). However, this scenario is not supported by the K-bentonite record: whereas

PRECORDILLERA K-BENTONITES the Precordilleran K-bentonites are Arenig to early Llanvirn in age, those associated with the Taconic o r o g e n y in the Appalachians are of m i d - C a r a d o c age. While the Precordillera m a y well have had its origins on the L a u r e n t i a n m a r g i n we do not find supporting evidence that it was still there during M i d - O r d o v i c i a n time (Astini et al. 1995; T h o m a s & Astini 1996). Rather, we find the ash pattern to be m o r e consistent with the palaeogeographical reconstructions of M a c Niocaill et al. (1997) who envisage drifting of the Precordillera in fairly close proximity to one or m o r e volcanic arcs with eventual collision along the A n d e a n m a r g i n of G o n d w a n a during the Ocloyic orogeny in MidOrdovician time. The extensive K - b e n t o n i t e record in the G u a n d a c o l region suggests that the source volcanoes m a y have been located n o r t h or northeast of the Precordillera. Conceivably, the P u n a - F a m a t i n a terrane could have been one of these volcanic arcs and might have served as one source of the K - b e n t o n i t e ashes, possibly in concert with active arc m a g m a t i s m on the G o n d w a n a plate itself, This research has been supported by NSF grants INT9513128 and EAR-9204893 to W.D.H, S.M.B, and D.R.K. We also acknowledge support from CONICET to C.C and R.A.A as part of a joint NSFCONICET agreement. The regional mining authority of San Juan Province provided field logistical support, and we were aided in our field work by A. Cuerda and D. Repol. Constructive reviews were provided by P. T. Leat, R. A. Batchelor, and C. W. Rapela.

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--

123

8r MARETTO, H. M. 1996. Anfilisis estratigr~fico del Silflrico de la Precordillera Central de San Juan y consideraciones sobre la evoluci6n de la cuenca. XIII Congreso Geol6gico Argentino y III Congreso de Exploraciones de Hidrocarburos, Buenos Aires, Actas, I, 351-368. , BENEDETTO, J. L. 8r VACCARI, N. E. 1995. The early Paleozoic evolution of the Argentine Precordillera as a Laurentian rifted, drifted, and collided terrane: A geodynamic model. Geological Society of America Bulletin, 107, 253-273. BAHLBURG,H. & HERVI~,F. 1997. Geodynamic evolution and tectonostratigraphic terranes of northwestern Argentina and northern Chile. Geological Society of America Bulletin, 109, 869-884. BENEDETTO, J. L. & SANCHEZ, T. M. 1996. Paleobiogeography of brachiopod and molluscan faunas along the South American margin of Gondwana during the Ordovician. In: BALDIS, B. A. & ACElqOLAZA,F. G. (eds) Early Paleozoic Evolution in NW Gondwana. Serie Correlacibn Geol6gica, 12. San Miguel de Tucum~m, 23-38. BERGSTROM, S. M., HUFF, W. D., KOLATA, D. R. & BAUERT, H. 1995. Nomenclature, stratigraphy, chemical fingerprinting, and areal distribution of some Middle Ordovician K-bentonites in Baltoscandia. Geologiska F6reningens I Stockholm F6rhandlingar, 117, 1-13. , --, KREKELER, M. P. S., CINGOLANI, C. ASTINI, R. A. 1996. Lower and Middle Ordovician K-bentonites in the Precordillera of Argentina: A progress report. XIII Congreso Geol6gico Argentino y H Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, V, 481-490. - - , YOST, D. A. & HART, C. 1997. A unique Middle Ordovician K-bentonite bed succession at R6stfinga, S. Sweden. Geologiska Fdreningens I Stockholm Fdrhandlingar, 119, 231-244. BOHOR, B. F. & TRIPLEHORN, D. M. 1993. Tonsteins." altered volcanic-ash in coal-bearing sequences. Geological Society of America Special Paper 285. BRUSSA, E. D. & ASTINI, R. A. 1997. La Biozona de Undulograptus austrodentatus en la Formacion Gualcamayo, Quebrada de Los Saltitos, suroeste de Guandacol, Precordillera Argentina. Reunion anual de Comunicaciones Paleontologicas de la Asociaci6n Paleontologica Argentina, Resumenes, Ameghiniana, 34, 113-127. CAREY, S. N., SIGURDSSON, H. & SPARKS, R. S. J. 1989. Experimental studies of particle-laden plumes. Journal of Geophysical Research, 93, 15314-15328. CINGOLANI, C. A., VARELA, R. & ACElqOLAZA,F. G. 1985. Caracteres geologicos y paleoicnologicos del Cerro del Corral (Partido de Necochea), Provincia de Buenos Aires. Comisi6n de Investigaciones Cientificas de la Provincia de Buenos Aires, Primeras Jornadas Geol6gicas Bonaerenses, 891-908. - - . , HUFF, W. D., BERGSTROM, S. M. & KOLATA, D. R. 1997. Bentonitas potfisicas Ordovicicas en la Precordillera de San Juan y su significaci6n tectomagmtica. Revista de la Asociaci6n Argentina, 52, 339-354.

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HUFF, W. D., BERGSTROM, S. M. & KOLATA, D. R. 1992. Gigantic Ordovician volcanic ash fall in North America and Europe: Biological, tectonomagmatic, and event-stratigraphic significance. Geology, 20, 875-878. , CINGOLANI,C. & DAVIS,D. W. 1995a. ]k,liddie Ordovician K-bentonites discovered in the Precordillera of Argentina: Geochemical and paleogeographic implications. In: COOPER, J. D., DROSER, M. L. & FINNEY, S. C. (eds) Ordovician Odyssey: Short papers for the Seventh International Symposium on the Ordovician System. The Pacific SEPM, Book 77, 343-349. - - 8,: SUN, H. 1998. The Lower Silurian ()smu'ndsberg K-bentonite. Part II: Mineralogy, geochemistry, chemostratigraphy, and tectonomagmatic significance. Geological Magazine, 135, 15-20. --, DAVIS, D. W., BERGSTROM, S. M., KREKELER, M. P. S., KOLATA,D. R. 8z CINGOLANI,C. 1997b. A biostratigraphically well-constrained K-bentonite U-Pb zircon age of the lowermost Darriwilian Stage (Middle Ordovician) from the Argentine Precordillera. Episodes, 20, 29-33. --, KOLATA, D. R., BERGSTROM, S. M. & CINGOLANI, C. 1995b. K-bentonite geochemistry and paleogeography as a means of testing the proposed justaposition of North and South America during the Ordovician. Geological Society of America, Abstracts with Program, 27, 63. & ZHANG, Y.-S. 1996. Largemagnitude Middle Ordovician volcanic ash falls in North America and Europe: dimensions, emplacement and post-emplacement characteristics. Journal of Volcanology and Geothermal Research, 73, 285-301. , CINGOLANI, C. t~ AST1NI, R. A. '1997a.' The Ordovician-Silurian evolution of the Iapetus Ocean: Paleogeographic implications from K-bentonite distribution patterns. Geological Society of America, Abstracts with Program, 2 9 , 281. , MORGAN, D. J. & RUNDLE, C. C. 1997c. Silurian K-bentonites of the Welsh Borderlands. Geochemistry, mineralogy and K-Ar ages of illitization. British Geological Survey, Technical Report WG/96/45. HI)NICKEN, M. A. & SARMIENTO,G. N. 1982. La Zone Baltoescandinava de Oepikodus evae (Conodonte, Arenigiano Inferior) en el perfil del Rio Guandacol, La Rioja, Argentina. Quinto Congreso Latinoamericano de Geologia, Actas, 1, 791-796. - & -1985. Oepikodus evae (Lower Arenigian Conodont) from Guandacol, La Rioja, Argentina. Boletin Academia, Nacional de Ciencias de C6rdoba, 56, 323-331. KELLER, M. & DICKERSON, P. W. 1996. The missing continent of Llanoria- was it the Argentine Precordillera? XIII Congreso Geol6gico Argentino y III Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, V, 355-367. KING, P. B. 1937. Geology of the Marathon Region, Texas. United States Geological Survey Professional Paper 187.

PRECORDILLERA K-BENTONITES KOLATA, D. R., FROST, J. K. & HUFF, W. D. 1987. Chemical correlation of K-bentonites in the Middle Ordovician Decorah Subgroup, upper Mississippi Valley. Geology, 15, 208-211. --, HUFF, W. O. & BERGSTROM, S. M. 1996. Ordovician K-bentonites of eastern North America. Geological Society of America Special Paper 313. KOUKHARSKY, M., COIRA, B., BARBER, E. & HANNING, M. 1988. Geoquimica de vulcanitas Ordovicicas de La Puna (Argentina) y sus implicancias tectonicas. 5 ~ Congreso Geologico Chileno, Santiago, Actas, III, 137-151. KREKELER, M. P. S., HUFF, W. D., KOLATA, D. R., BERGSTROM, S. M. & CINGOLANI, C. 1995. Mineralogy and grain characteristics of Middle Ordovician K-bentonites from the Precordillera of Argentina. In: COOPER, J. D., DROSER, M. t . & FINNEY, S. C. (eds) Ordovician Odyssey: Short Papers for the Seventh International Symposium on the Ordovician System. The Pacific Section SEPM, Book 77, 355-356. LAURENCE, R. A. 1944. An Early Ordovician sinkhole deposit of volcanic ash and fossiliferous sediments in East Tennessee. Journal of Geology, 52, 235-249. LEHNERT, O. 1995a. Ordovizische Conodonten aus der Pr~ikordillere Westargentiniens: ihre Bedeutung fiir Stratigraphie und Palfiogeographie. Erlanger Geologische Abhandlungen, 125, 1-193. - - 1 9 9 5 a . The Tremadoc/Arenig transition in the Argentine Precordillera. In: COOPER, J. D., DROSER, M. L. & FINNEY, S. C. (eds) Ordovician Odyssey: Short Papers for the Seventh International Symposium on the Ordovician System. The Pacific Section SEPM, Book 77, 145-148. & KELLER, M. 1994. The conodont record of the Argentine Precordillera: Problems and possibilities. Zentralblatt ffir Geologie und Paldontologie 1993, 231-244. LOPEZ DE LUCHI, M. G. & DALLA SALDA, L. H. 1995. Some Pampean PreTaconic and Famatinian Taconic Granites of Southwestern South America. In: BROWN, M. (ed.) Third Hutton Symposium on Granites and related rocks. proceedings of a symposium heM at the University of Maryland at College Park, 27 August to 2 September 1995. Geological Society of America Special Papers, 315, 1-43. MAC NIOCAILL, C., VANDER PLUIJM, B. m. 8z VANDER VOO, R. 1997. Ordovician paleogeography and the evolution of the Iapetus ocean. Geology, 25, 159-162. MANNHEIM, R. 1993. Genesis de las volcanitas Eopaleozoicas del Sistema de Famatina, noroeste de Argentina. XII Congreso Geol6gico Argentino y H Congreso de Exploracidn de Hidrocarburos, Mendoza, Actas, IV, 147-155. MERRIMAN, R. J. & ROBERTS, B. 1990. Metabentonites in the Moffat Shale Group, Southern Uplands of Scotland: Geochemical evidence of ensialic marginal basin volcanism. Geological Magazine, 127, 259-271. NASH, W. P. 1992. Analysis of oxygen with the electron microprobe: Applications to hydrated glass and minerals. American Mineralogist, 77, 453-457. -

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ORTEGA, G., ALBANESI, G., OTTONE, E. & HOLFELZ, G. 1996. Conodonts, palynomorphs and graptolites from the Los Azules Formation, Argentine Precordillera. Sixth International Conodont Symposium (ECOS V1), Abstracts, 42. PANKHURST, R. J., RAPELA, C. W., SAAVEDRA, J., BALDO, E., DAHLQUIST, J., PASCUA, I. & FANNING, C. M. The Famatinian Magmatic arc in the central Sierras Pampenas: an Early to MidOrdovicean Continental arc on the Gondwana margin. This volume. PEARCE, J. A., HARRIS, N. B. W. & TINDLE, A. G. 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of Petrology, 25, 956-983. PROKOPENKO, M., KREKELER,M. P. S., HUFF, W. D., BERGSTROM,S. M., KOLATA,D. R., CINGOLANI,C. LEHNERT, O. T. H. 1997. Ordovician K-bentonites in the Argentine Precordillera: origin, composition, diagenesis, and paleoenvironmental implications. The llth International Clay Conference, Program with Abstracts, 60. RAMOS, V. A., JORDAN, T. E., ALLMENDINGER,R. W., MPODOZIS, C., KAY, S. M., CORTI~S, J. M. & PALMA, M. 1986. Paleozoic terranes of the central Argentine-Chilean Andes. Tectonics, 5, 855-880. RAPELA, C. W., TOSELLI, A., HEAMAN, L. & SAAVEDRA, J. 1990. Granite plutonism of the Sierras Pampeanas; an inner Cordilleran Paleozoic arc in the southern Andes. In: KAY, S. M. & RAPELA, C. W. (eds) Plutonism from Antarctica to Alaska. Geological Society of America Special Papers, 241, 77-90. - - , COIRA, B., TOSELLI,A. & SAAVEDRA,J. 1992. E1 magmatismo del Paleozoico inferior en el Sudoeste de Gondwana. In: GUTIERREZ MARCO, J. C., SAAVEDRA, J. & RABANO, J. (eds) Paleozoico Inferior de Ibero-America. Universidad de Extremadura, Spain, 21-68. REYNOLDS, R. C. J. 1985. NEWMOD: A computer program for the calculation of one-dimensional diffraction patterns of mixed-layer clays. R. C. Reynolds, 8 Brook Rd., Hanover, NH, USA. ROSE, W. I. & CHESNER, C. A. 1987. Dispersal of ash in the great Toba eruption, 75 ka. Geology, 15, 913-917. Ross, R. J. J., ADLER, F. J., AMSDEN,T. W. et al. 1982. The Ordovician System in the United States: Correlation Chart and Explanatory Notes. International Union of Geological Sciences, Publication No. 12. SACK, R. O. & GHIURSO, M. S. 1994. Thermodynamics of multicomponent pyroxenes: I. Formulation of a general model. Contribution to Mineralogy and Petrology, 116, 277-286. SIGURDSSON, H. & CAREY, S. 1989. Plinian and coignimbrite tephra fall from the 1815 eruption of Tambora volcano. Bulletin of Folcanology, 51, 243-270. SPRAY, J. G. t~ RAE, D. A. 1995. Quantitative electron-microprobe analysis of alkali silicate glasses: a review and user guide. The Canadian Mineralogist, 33, 323-332.

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The geochemistry and provenance of Ordovician turbidites in the Argentine Puna HEINRICH

BAHLBURG

Geologisch-Palgiontologisches Institut, Westfdilische Wilhelms-Universitgit, C o r r e n s s t r a s s e 24, 48149 M f i n s t e r , G e r m a n y

Abstract: Petrographical and major, trace and rare earth element data are combined in a provenance study of Early and Mid-Ordovician volcaniclastic turbidites in the Puna of NW Argentina. The tectonic setting of the depositional basin behind a west-facing volcanic arc is well established from the literature, so that the new data may be used to test published geochemical discrimination schemes of tectonic setting of sedimentary rocks. However, there is a controversy whether volcanism took place in an island arc or a mature continentalmargin arc. The data show that the arc had a sediment source of average upper crustal composition dominated by silicic and intermediate magmatic rocks. Major elements document that subaerial weathering significantly influenced the detritus composition, whereas the HFS elements in particular retained the original source signal. The HFS and REE characteristics, as well as other parameters, compare favourably with those of Recent deep sea sands derived from continental arcs.

The mineralogical and geochemical composition of sedimentary rocks ideally reflects the nature of the source regions, and indirectly the tectonic setting of the basin where the erosional debris is deposited (e.g., Dickinson & Suzcek 1979; Valloni & Mezzardi 1984; Bhatia & Crook 1986; McLennan et al. 1993). However, even in first cycle sands the interpretation of provenance and tectonic setting has to consider the effects of derivation from several distinct sources, weathering in the source regions, sorting of mineral grains during transport, and diagenesis (Nesbitt & Young 1984, 1989; Johnsson 1993; McLennan et al. 1993). In combination with evidence from petrography, geochemical analysis of major, trace and rare earth elements may give information on the extent of these effects, as well as on the presence of components not readily recognisable under the microscope. As long as the bulk composition of a rock is not altered, geochemical analysis is a valuable tool also in the study of matrix rich sandstones (McLennan et al. 1993). This technique has been applied successfully to the interpretation of provenance and tectonic setting of sedimentary successions (e.g., Bhatia 1983; Roser & Korsch 1986; McLennan et al. 1990; Floyd et al. 1991; Roser et al. 1993), including cases of terranes of unknown derivation (Mortimer & Roser 1992). The Ordovician basin in the Puna of N W Argentina evolved to the east (present coordinates) of a coeval and extended arc system connected to eastward subduction beneath South America. The arc trended roughly N N W - S S E from northern Chile at c. 20~ into Argentina, where it can be traced at least to 32~ (Palma

et al. 1986; Bahlburg 1990; Rapela et al. 1992; Omarini & Sureda 1993; Pankhurst et al. this volume). There is continuing discussion about (i) whether the basin behind the arc north of 27~ was at least partially floored by oceanic crust or by stretched continental crust (Allmendinger et al. 1983; Bahlburg 1990; Forsythe et al. 1993; Blasco et al. 1996; Bahlburg & Herv6 1997) and (ii) whether the volcanic source was an island arc or a continental-margin arc (Rapela et al. 1992; Breitkreuz et al. 1989; Pankhurst et al. this volume). Sedimentary rocks which were deposited contemporaneously with arc activity are well exposed and cover large areas in the N W Argentine Puna between 22~ and 26~ Because the intensity of Late Ordovician folding and the Cenozoic tectonic overprint increases southwards, this study concentrates on the Ordovician deposits in the northern Puna between 23~ and c. 24~ In this paper I will present the first whole-rock geochemical data set on Ordovician turbidite sandstones from the Argentinian Puna (Table 1). In combination with a summary of published petrographical data, they will be used to extract more detailed information on the evolution of the sources of the turbidite deposits, i.e., island arc versus continental margin arc, and on the extent of subaerial weathering in the source terrains. Since the tectonic setting of the basin, which evolved from back-arc to retro-arc during the Ordovician, is relatively well established (Bahlburg 1991; Rapela et al. 1992; Bahlburg & Herv6 1997), existing plate-tectonic discrimination schemes for sedimentary rocks can also be tested.

BAHLBURG,H. 1998. The geochemistry and provenance of Ordovician turbidites in the Argentine Puna. In: PANKHURST, R. J. & RAPELA,C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 127-142.

128

H. BAHLBURG

Table 1. Synopsis of selected framework mineral data o f modern deep-sea turbidite sands and o f the Ordovician Puna turbidites Basin

Q

F

L

P/F

Lv/L

86.8 73.1 47.7

9.4 26.1 49.8

3.8 0.7 2.5

0.44 0.34

-

45.0 64.0 54.4

32.3 33.6 35.6

22.6 2.4 I0.0

0.40 0.27 0.34

0.01 -

5.0 33.3 54.8 2.7

9.0 49.0 35.4 19.7

86.0 17.7 9.8 77.6

1.00 0.79 0.47 0.96

1.00 1.00 1.00

67.0 49.0 49.0

8.0 23.0 17.0

25.0 28.0 34.0

0.72 0.64 0.81

0.62 0.85 0.98

Passive margins

Bay of Biscay Guinea South Korea Continental collision

Ganges Ganges Ganges Continental arc

Peru-Chile Peru-Chile Java Java Ordovician Puna basin

UTS LTS VS

Sources: Valloni & Maynard (1981), Maynard (1984), Yerino & Maynard (1984), compilation by McLennan et al. (1990), and Bahlburg (1990).

Stratigraphy and geological setting In the Early and early Mid-Ordovician two wide marine facies belts covered NW Argentina: a shallow-marine siliciclastic platform in the east, i.e., the present-day Cordillera Oriental, and a deeper basin to the west in the Puna highland (Fig. 1; Turner 1970; Salfity et al. 1975; Gohrbrandt, 1992; Moya et al. 1994). In the east, these rocks overlie Late Cambrian quartz sandstones and Late Proterozoic-Early Cambrian folded metaturbidites; in the west, the base of the Ordovician is not exposed. During the Llanvirn, emergence of the Cordillera Oriental region led to the separation of the eastern platform from the western Puna basin (Salfity et al. 1984; Ornarini & Sureda 1993). The Ordovician deposits of the Puna are the object of this study and can be grouped into two lithological associations: (i) the c. 4000m thick Volcanosedimentary Successions of Early Ordovician age, and (ii) the Puna Turbidite Complex of Mid- to early Late Ordovician age. The Puna Turbidite Complex consists of the Lower Turbidite System (Llanvirn, c. 2000m thick), and the Upper Turbidite System (Fig. 1; Llandeilo-early Caradoc, c. 900m thick, Bahlburg et al. 1990; Moya et al. 1993). The Volcanosedimentary Successions formed as a volcaniclastic apron on the eastern flank of the volcanic arc. They consist of basalt and andesite lavas, mafic hydroclastic breccias, dacites and rhyolites, which are associated with

silicic lapilli tufts and thick successions of volcaniclastic turbidites (Breitkreuz 1986; Coira & Barber 1987; Koukharsky et al. 1988; Breitkreuz et al. 1989; Bahlburg 1990). The volcanic rocks are calc-alkaline and of volcanic arc affinity (Koukharsky et al. 1988; Breitkreuz et al. 1989; Bahlburg 1990; Rapela et al. 1992). The uppermost Arenig strata interdigitate with and are overlain by the Lower and Upper Turbidite Systems, respectively. Detritus was supplied from the arc, eastwards to the basin, through several channel systems. Palaeocurrents indicated by flute marks were directed mostly northwards in the Early Ordovician and Llanvirn, and southwards in the Llandeilo-early Caradoc (Bahlburg 1991). Tectonic subsidence related to the progressive construction of the Ordovician arc increased markedly from the Arenig to the Llanvirn and then decreased (Bahlburg & Furlong 1996). The basin fill was folded after the early Caradoc and before the end of the Ordovician in the Ocl6yic orogeny (Turner 1970).

General sample properties Distribution and normal grading are present in the turbidites of all units. When possible samples were taken from the coarse sand, basal turbidite divisions Ta or Tb. As shales sample widespread crustal regions, they tend to reflect the average crustal composition (Taylor & McLennan 1985). Therefore, in sampling the coarser layers of the

ORDOVICIAN TURBIDITES IN THE PUNA

129

Fig. 1. Outcrop map of Ordovician clastic rocks in the Puna of northwest Argentina. Section indices correspond to sample numbers in Table 2. deposit the least homogenised part was obtained. Clasts show all grades of rounding, with a high proportion of angular clasts. Euhedral feldspar is common but not abundant. The analysed samples are carbonate-free and mostly mediumgrained sandstones with an average matrix content of 7% in the Volcanosedimentary Successions, and of 9% and 13%, respectively, in the Lower and Upper Turbidite Systems. The clay mineral spectrum consists mainly of illite and chlorite, smectite is present to 4-5% in some samples and indicates that diagenetic to

low anchizonal metamorphic conditions were not exceeded. Pressure-solution cleavage and matrix recrystallization were not observed.

Petrographical characteristics The turbidite sandstones of the Volcanosedimentary Successions are quartz intermediate with 35-60% quartz (Crook 1974; Bahlburg 1990). The rocks are poorly sorted and clasts of all sand sizes occur. Monocrystalline quartz is

130

H. BAHLBURG

Fig. 2. Framework modes of Ordovician turbidite sandstones from the Puna of NW Argentina (data: Bahlburg 1990). VS, Volcanosedimentary Successions; LTS, Lower Turbidite system; UTS, Upper Turbidite system. Q, total quartz; F, total feldspar; L, total lithics; Qm, monocrystalline quartz; Lt, total lithics plus chert; Qp, polycrystalline quartz; Lv, volcanic rock fragments; Ls, sedimentary rock fragments. almost exclusively embayed and magmatic in origin. Embayments are frequently filled with very fine-grained polycrystalline quartz representing the groundmass of silicic magmatic rocks. Feldspar (plag > KF) is around 25% but may represent up to 35% of the grains, andesine is the most An-rich member observed; mica (muscovite) is rare. Glass shards are abundant in some layers, alterd to either chlorite or chalcedony. Volcanic clasts vary between 25 and 50% and consist mostly of basalt and andesite fragments; sedimentary clasts were not observed. The sandstones of the turbidite systems have the same general characteristics. They are also quartz intermediate with the Upper Turbidite System slightly more quartz-rich. Clasts of (meta-)sedimentary rocks are present and increase from 7% in the Lower to a maximum of 15% in the Upper Turbidite System. Gneiss fragments are rare and represent the highest metamorphic grade encountered in the detritus. Muscovite is locally abundant, biotite is rare. Glass shards were not found. Zircon, apatite, hornblende, tourmaline and pyroxene occur with decreasing abundance. The framework modes of the turbidite sandstones obtained through point counting (Dickinson & Suczek 1979; Dickinson 1985) reflect the overall composition. The average framework mode of the Volcanosedimentary Successions is Q49F17L34, and Q49F23L2s for the Lower and Q67FaL25 for the Upper Turbidite System,

respectively (Fig. 2; Table 1). In terms of the discrimination diagrams of Dickinson (1985) the QFL and QmFLt modes classify the rocks as recycled orogenic material. This field includes compositions derived from collision orogens, foreland uplifts and recycled accretionary wedges (Dickinson & Suczek 1979). In the QpLvLs diagram the data points do not fall into any of the discriminating fields. Sedimentology and the regional context show that the rocks were derived from a relatively evolved and partially subaerial volcanic arc with abundant intermediate and silicic rocks (e.g., Koukharsky et al. 1988; Breitkreuz et al. 1989; Rapela et al. 1992). This is reflected in the detritus by the abundance of mono- and poly-crystalline quartz of magmatic origin. This abundance probably biases the framework modes towards the Q-apices of the diagrams, thus masking a potential apparent arc origin of the detritus. Petrographical features indicate that the detritus spectrum in question is relatively rich in labile minerals and rock fragments. Sediment reworking in intermittent storage areas has apparently not been efficient. The abundance of coarse- and medium-grained dense minerals and rock fragments indicates that sorting, although always present in turbidites, was minor. Framework modes of the Ordovician Puna samples are similar to modern deep sea turbidite sands from both continental collision basins and continental arc basins, such as the Peru-Chile or Java

ORDOVICIAN TURBIDITES IN THE PUNA trenches (Table 1). In combination with ratios of plagioclase to total feldspar (P/F) and volcanic fragments to total fragments (Lv/L) the data show more affinity to detritus accumulated in continental arc basins than to detritus in continental collision basins (Table 1).

Geochemical features Weathering and sediment recycl&g

According to their SiO2 contents, most of the analysed sandstones are quartz-intermediate and quartz-rich sensu Crook (1974; Tables 2-4, Fig. 3). The use of immobile trace element ratios Nb/Y and Zr/TiO2 as a proxy for bulk geochemistry (Floyd & Winchester 1975) shows that all three units have a dacitic to rhyodacitic composition, with four samples having characteristics of andesites. The chemical index of alteration (CIA, Nesbitt & Young 1982) is an indicator for bulk chemical changes due to weathering and the conversion of feldspars, volcanic glass and other labile components to clay minerals. The feldspar-rich sandstones QJ 9a, QJ 12 and HY 19-1 (Table 1) have CIA values of c. 57, 60 and 64, which are close to values of c. 50 characteristic of fresh granites and rhyolites (Fig. 3). CIA values go up to 78, but the average for all samples is 70 (Fig. 3). Igneous lithologies in the source areas include rhyolites, granites, granodiorites, andesites and basalts, the last of which would be expected to have CIA values of between 30 and 45. Illite, an abundant diagenetic and low grade metamorphic product in the samples, has nominal values of between 75 and 85. Values of strongly weathered residual clays reach 100 (Nesbitt & Young 1982). From the Volcanosedimentary Successions through the 90,00 80,00 nm

niini~,~, .',

9

70,00~oo

60,00

o

R~o

~ o o o

50,00 40,00 40,00

50100

60100 70100 80100 90100 SiO 2 (wt %)

Fig. 3. Plot of SiO2 v. CIA (chemical index of alteration, Nesbitt & Young 1982). CIA = A1203/[100 x (AI20 3 + K20 + Na20 + CaO*)], CaO* calculated carbonate free.

131

Lower to the Upper Turbidite System the CIA values increase from a minimum of 57 to a maximum of 78 (Fig. 3), indicating that the detritus feeding the turbidite systems underwent progressively more weathering through time in the source area or en route to the basin, respectively. Reworking and sorting during sediment transfer leads to the concentration in the deposits of some resistent heavy minerals, commonly zircon. Geochemically this leads to an increase in Zr concentration, which therefore serves as an indicator of reworking. When combined with Th and Sc which are, respectively, incompatible and compatible in igneous differentiation processes, the Zr/Sc ratio is an indicator of zircon abundance and thus the degree of reworking, whereas Th/Sc relates to the degree of igneous chemical differentiation (Fig. 4; McLennan et al. 1993). The Ordovician Puna sedimentary rocks have relatively elevated Th/Sc ratios and moderate Zr/Sc. This is taken to reflect the absence of significant reworking and the input from evolved igneous sources the composition of which largely corresponds to the average upper crust.

Provenance

The petrographical composition of the turbidites, the geochemical features of the intercalated mafic and silicic lavas and the regional geological context, all suggest that the studied rocks were derived from a volcanic arc (Schwab 1973; Koukharsky et al. 1988; Breitkreuz et al. 1989; Rapela et al. 1992). Regarding the arc along the western edge of the Puna, some authors interpret it as an island arc (e.g., Rapela et al. 1992), whereas others consider it a continental margin arc residing on continental crust (Breitkreuz et al. 1989). The evidence supplied by the geochemical data on the arc-derived turbidites offers some arguments to resolve this discrepancy. The thickness and volume of exposed mafic rocks are considerably less than those of the silicic rocks. Although the LIL elements K and Rb are considered to be relatively mobile during diagenesis and low-grade metamorphism, the K/Rb ratio in the samples is compatible with derivation of the turbidites from mainly silicic magmatic precursors (Fig. 5), a case also observed in Rhenohercynian greywackes from Germany and Great Britain (Floyd & Leveridge 1987; Floyd et al. 1990). In contrast, most HFS elements are immobile and are quantitatively transported from source to sink, so that they are of particular value in identifying provenance (Taylor & McLennan 1985). The upper continental crust has a Th/Sc ratio of > 1 (Taylor

Table 2. Geochemical data for the Ordovician volcanosedimentary successions

Sample

FP l

FP 15 73.86 13.38 3.13 0.85 0.06 0.91 4.65 1.54 0.47 0.12 1.77 100.74

QJ 9a 54.35 16.81 7.92 3.93 0.13 4.83 5.18 2.60 0.99 0.65 2.32 99.70

QJ 12

QJ 13"

QJ 15"

AP 12

AP 22

61.37 14.13 6.04 4.27 0.12 3.89 2.88 2.32 0.66 0.17 3.86 99.71

66.69 13.11 5.00 3.42 0.06 3.00 1.73 1.58 0.48 0.20 4.15 99.42

81.05 9.64 1.76 0.60 0.00 0.26 2.08 1.54 0.27 0.10 2.56 99.86

81.34 10.24 0.99 0.19 0.01 0.67 4.50 1.09 0.15 0.04 0.73 99.95

78.14 11.64 0.32 0.19 0.00 0.15 1.20 6.48 0.09 0.02 0.77 99.01

75.38 12.72 1.74 0.51 0.03 0.37 5.01 1.64 0.25 0.05 1.63 99.34

67.90 14.12 4.81 1.26 0.08 1.36 4.98 1.39 0.50 0.08 3.24 99.72

25.0 42.0 32.0

22.0 37.0 29.0

61.0 18.0 206.0 103.0 170.0 1 0 5 0 . 0 1 3 3 9 . 0 444.0 6.0 85.0 71.0 98.0 1.0 7.0 10.0 8.0 3.0 3.0 3.0 9.0 11.0 8.0 14.0 19.0 42.0 20.0 26.0 25.0 6.0 6.0 7.0 18.0 7.0 33.0 1.0 15.0 14.0 2.0 8.0 4.0 4.0 1.0 2.0 0.0 0.0 3.0 16.0 38.0 23.0 6.0 9.0 11.0 1.0 2.0 1.0

46.0 31.0 158.0 461.0 3.0 160.0 3.0 9.0 2.0 7.0 11.0 27.0 11.0 79.0 21.0 12.0 2.0 2.0 57.0 17.0 1.0

SiO2 A1203 Fe203 MgO MnO CaO Na20 K20 TiO2 P205 LOI Total

76.91 13.65 0.61 0.38 0.01 0.54 3.27 1.78 0.36 0.09 1.70 99.31

La Ce Nd Sm Eu Gd Dy Er Yb Lu

49.0 51.0 39.0

28.0 42.0 30.0

124.0 214.0 103.0

35.0 55.0 25.0

28.9 51.5 23.3 4.6 1.0 4.2 3.7 2.0 2.0 0.4

27.1 49.1 24.3 5.2 0.9 5.1 4.2 2.3 2.2 0.4

25.0 29.0 15.0

16.0 31.0 23.0

Rb Cs Sr Ba Hf Zr Ta Nb U Th Pb Y Sc V Cr Co Ni Cu Zn Ga Sn Total REE

71.0

58.0

55.0

61.0

52.0

56.0

39.0

194.0

283.0 322.0

237.0 1104.0

285.0 522.0

124.0 327.0

132.0

139.0

292.0

118.0

76.0

110.0

9.0 2.0 9.0 9.0 28.0 9.0 50.0 9.0 1.0 2.0 0.0 16.0 11.0 2.0

11.0 2.0 10.0 14.0 25.0 10.0 70.0 16.0 8.0 6.0 4.0 50.0 12.0 1.0

40.0 8.0 29.0 23.0 31.0 19.0 186.0 93.0 21.0 39.0 49.0 89.0 18.0 0.0

10.0 3.0 8.0 13.0 25.0 18.0 140.0 115.0 16.0 38.0 32.0 54.0 15.0 0.0

6.0

6.0

9.0

10.0

19.0 17.0 155.0 89.0 17.0 46.0 62.0 66.0 10.0 0.0 121.49

24.0 7.6 35.0 20.0 10.0 16.0 11.0 47.0 0.0 0.0 120.81

14.50 5.70 3.95 1.71 0.54

12.13 8.83 3.29 1.87 0.40

1.70 0.53

3.56 1.32

5.09 0.91 1.09 0.18

8.41 0.74 1.35 0.45

668.0 428.0 1550.0 8 1 5 . 0

La/Yb LaN/YbN LaN/SmN GdN/YbN Eu/Eu* La/Sc Th/Sc Th/U SIO2/A1203 K20/Na20 Na20/K20 Rb/Sr

5.44 1.00 4.50 5.64 0.54 1.84 0.25

2.80 1.00 5.00 5.52 0.33 3.02 0.24

6.53 1.53 3.63 3.23 0.50 1.99 0.08

1.94 0.44 2.67 4.34 0.81 1.24 0.14

4.17 1.50 3.00 7.94 0.24 4.13 0.19

2.67 1.83 3.67 6.71 5.40 0.19 1.88

HY 13

3.57 1.14 2.67 5.93 0.33 3.05 0.36

HY 19

2.00 0.64 3.50 4.81 0.28 3.57 0.29

XFA whole rock analyses of major and trace elements were conducted at the Institut ffir Mineralogie, Freie Universitfit Berlin and the Centre de Recherches P~trographiques et G+ochimiques (CRPG), Nancy. REE concentrations were determined at the CRPG by optical emission spectrochemical analysis using microwave plasma excitation (Govindaraju et al. 1976). Samples were crushed in an agate mill.

H Y 19-1

H Y 20

H Y 21

H Y 22

H Y 23

H Y 23b

66.84 14.38 5.16 1.27 0.07 1.46 5.33 1.43 0.52 0.17 3.69 100.32

77.30 11.14 2.66 0.69 0.03 0.63 2.49 2.28 0.44 0.08 1.77 99.51

81.91 8.40 1.51 0.35 0.03 1.81 3.22 0.63 0.25 0.05 2.08 100.24

84.12 7.67 1.59 0.30 0.03 0.29 3.28 0.34 0.37 0.08 0.97 99.03

74.58 13.74 1.83 0.28 0.02 0.30 6.96 0.25 0.41 0.08 0.93 99.38

78.51 10.68 1.11 0.42 0.02 0.88 3.20 1.84 0.24 0.05 2.41 99.37

80.53 9.00 2.54 0.51 0.00 0.35 2.08 1.79 0.34 0.14 1.94 99.22

68.93 14.38 5.01 1.97 0.06 1.09 2.17 4.07 0.67 0.19 1.26 99.78

70.04 11.36 2.71 1.08 0.04 1.20 3.45 1.70 0.38 0.12 2.03

18.64 3.62 2.14 1.30 0.04 1.33 1.69 1.35 0.22 0.14 1.12

19.4 44.7 16.7 4.2 1.1 3.8 3.7 2.2 2.3 0.7

29.0 48.0 32.0

15.0 32.0 26.0

20.0 36.0 26.0

33.0 36.0 28.0

24.0 36.0 28.0

20.9 40.2 16.0 3.5 0.8 3.0 2.8 1.6 1.6 0.3

23.0 36.0 25.0

30.1 48.6 28.7 3.5 0.7 3.2 2.9 1.6 1.6 0.4

24.7 41.9 19.8 1.8 0.4 1.8 1.5 0.9 0.8 0.2

44.0

78.0 6.0 116.0 497.0 9.0 211.0 1.0 10.0 3.0 8.0 9.0 27.0 8.0 47.0 20.0 5.0 6.0 3.0 42.0 12.0 1.0

22.0 8.0 158.0 226.0 5.0 107.0 1.0 7.0 3.0 10.0 14.0 24.0 6.0 26.0 12.0 4.0 1.0 1.0 30.0 8.0 1.0

16.0 24.0 123.0 210.0 5.0 76.0 1.0 7.0 1.0 6.0 13.0 23.0 8.0 30.0 18.0 4.0 1.0 3.0 36.0 7.0 2.0

10.0 12.0 97.0 154.0 5.0 117.0 1.0 8.0 3.0 12.0 12.0 23.0 9.0 58.0 20.0 4.0 2.0 1.0 45.0 10.0 2.0

61.0 11.0 97.0 561.0 7.0 154.0 1.0 8.0 1.0 7.0 8.0 24.0 7.0 46.0 19.0 5.0 5.0 2.0 28.0 9.0 1.0

55.0

57.0 11.0 207.0 548.0 5.0 89.0 3.0 7.0 1.0 7.0 11.0 22.0 6.0 36.0 11.0 3.0 2.0 1.0 40.0 9.0 2.0

54.4 13.8 196.3 594.2 5.0 126.9 1.1 9.6 2.6 8.8 14.4 23.1 9.7 61.7 29.1 8.0 10.3 11.1 41.2 9.6 0.8

39.7 9.4 147.1 410.3 2.5 61.3 0.8 7.7 1.8 5.8 9.2 6.3 5.0 49.7 32.2 5.8 14.3 17.6 20.7 4.7 0.8

176.0 484.0 171.0 9.0

24.9 18.0 83.0 20.0 15.0 6.0 13.0 66.0 12.0 0.0 8.43 98.80 5.70 2.91 1.34 0.67 1.08 0.00 4.65 0.27 3.73 0.25

3.63 1.00 2.67 6.94 0.92 1.09 0.67

2.50 1.67 3.33 9.76 0.19 5.14 0.14

2.50 0.75 6.00 10.96 0.10 9.74 0.13

3.67 1.33 4.00 5.43 0.04 28.12 0.10

3.43 1.00 7.00 7.35 0.57 1.74 0.63

H Y 23b-1 H Y 29

105.0 629.0 167.0 8.0 6.0 19.7 7.9 47.0 21.0 8.0 8.0 14.0 38.0 6.0 0.0 13.06 90.70 8.83 3.76 1.52 0.60 2.65 0.76 8.95 0.86 1.16 0.52

3.83 1.17 7.00 4.79 1.87 0.53 0.28

Mean

3.03 0.98 3.91 6.13 0.78 3.83 0.36

St. dev.

1.44 0.50 1.76 2.43 1.17 6.13 0.40

134

H. B A H L B U R G

Table 3. Geochemical data for the Lower Turbidite System Sample

MC3

MC5

MC6

MC9

MC 15

AH5

AH 12

AH 14

AH 18

AH22

SiO2 A1203 Fe203 MgO MnO CaO Na20 K20 TiO2 P205 LOI Total

78.50 10.73 1.28 0.60 0.01 0.09 2.57 1.67 0.41 0.06 3.95 99.86

75.32 11.94 4.21 1.07 0.02 0.14 1.40 2.53 0.44 0.07 2.49 99.62

84.65 6.78 2.80 0.74 0.02 0.06 1.11 1.07 0.33 0.04 1.49 99.09

71.31 14.74 3.93 1.18 0.02 0.04 1.50 3.49 0.53 0.10 2.80 99.62

79.71 10.68 2.08 0.61 0.01 0.14 3.15 1.26 0.45 0.07 1.22 99.39

75.06 11.86 3.84 0.88 0.03 0.57 2.21 2.29 0.54 0.11 2.22 99.60

78.36 10.26 2.84 0.96 0.03 0.13 2.34 2.54 0.38 0.08 1.48 99.39

62.69 19.33 6.75 2.42 0.05 0.13 0.52 5.05 0.77 0.11 1.83 99.63

76.25 11.35 3.40 0.95 0.03 0.17 2.38 2.33 0.39 0.08 1.83 99.15

76.63 10.52 4.38 1.21 0.04 0.25 1.58 1.90 0.64 0.12 2.06 99.33

La Ce Nd Sm Eu Gd Dy Er Yb Lu

44.0 72.0 42.0

22.0 28.0 23.0

20.0 24.0 23.0

30.0 47.0 30.0

26.0 43.0 29.0

33.0 48.0 34.0

31.0 47.0 20.0

29.0 46.0 39.0

32.0 36.0 35.0

30.0 45.0 22.0

Rb Cs Sr Ba Hf Zr Ta Nb U Th Pb Y Sc V Cr Co Ni Cu Zn Ga Sn Total REE

73.0

113.0

48.0

156.0

59.0

101.0

85.0

214.0

87.0

78.0

107.0 208.0

67.0 325.0

48.0 205.0

91.0 449.0

168.0 203.0

75.0 392.0

83.0 618.0

48.0 637.0

104.0 402.0

54.0 282.0

194.0

240.0

237.0

164.0

177.0

177.0

143.0

119.0

140.0

210.0

11.0 3.0 11.0 16.0 32.0 9.0 44.0 19.0 2.0 3.0 0.0 12.0 11.0 3.0

11.0 2.0 11.0 18.0 27.0 8.0 51.0 17.0 7.0 10.0 10.0 66.0 14.0 2.0

8.0 2.0 7.0 18.0 21.0 6.0 27.0 11.0 4.0 6.0 3.0 32.0 7.0 1.0

12.0 3.0 12.0 22.0 34.0 11.0 67.0 26.0 6.0 10.0 1.0 46.0 17.0 3.0

9.0 2.0 10.0 16.0 26.0 8.0 43.0 14.0 4.0 4.0 0.0 34.0 8.0 1.0

11.0 2.0 11.0 23.0 28.0 9.0 65.0 27.0 7.0 15.0 3.0 72.0 14.0 4.0

9.0 4.0 9.0 18.0 26.0 7.0 50.0 18.0 5.0 10.0 0.0 43.0 10.0 2.0

16.0 2.0 14.0 17.0 35.0 15.0 116.0 73.0 12.0 15.0 14.0 68.0 26.0 4.0

10.0 3.0 9.0 21.0 27.0 8.0 45.0 19.0 5.0 12.0 0.0 60.0 11.0 2.0

11.0 3.0 10.0 16.0 26.0 9.0 72.0 30.0 8.0 17.0 5.0 61.0 13.0 3.0

La/Yb LaN/YbN LaN/SmN GdN/YbN Eu/Eu* La/Sc Th/Sc Th/U SiOz/A1203 K20/Na20 Na20/K20 Rb/Sr

4.89 1.22 3.67 7.32 0.65 1.53 0.68

Comments as for Table 2.

2.75 1.38 5.50 6.31 1.81 0.55 1.69

3.33 1.17 3.50 12.49 0.96 1.04 1.00

2.73 1.09 4.00 4.84 2.33 0.43 1.71

3.25 1.25 5.00 7.46 0.40 2.49 0.35

3.67 1.22 5.50 6.33 1.04 0.97 1.35

4.43 1.29 2.25 7.64 1.09 0.92 1.02

1.93 0.93 7.00 3.24 9.70 0.10 4.46

4.00 1.13 3.00 6.72 0.98 1.02 0.84

3.33 1.11 3.33 7.28 1.20 0.83 1.44

O R D O V I C I A N T U R B I D I T E S IN T H E P U N A

OL 3

135

OL 5*

OL 12"

OL 15"

OL 17

OL 18

OL 21

OL 26

74.98 12.72 3.57 1.04 0.02 0.19 1.55 2.42 0.48 0.12 2.10 99.18

74.00 12.67 3.24 1.11 0.02 0.12 2.04 3.20 0.40 0.13 2.69 99.62

76.34 10.44 3.09 0.83 0.02 0.17 1.92 3.17 0.35 0.13 2.09 98.55

74.60 11.53 3.12 0.88 0.02 0.25 2.16 2.50 0.38 0.11 3.07 98.62

73.55 12.22 4.15 1.43 0.05 0.32 1.88 2.53 0.69 0.14 2.23 99.20

76.49 11.20 2.94 1.00 0.04 0.30 2.36 2.64 0.38 0.08 1.69 99.13

76.47 10.52 3.51 1.12 0.05 0.42 2.06 2.84 0.45 0.11 1.65 99.20

70.10 13.10 7.25 1.81 0.06 0.29 1.36 2.38 0.71 0.17 2.79 100.02

32.0 52.0 29.0

28.6 58.0 25.7 5.5 1.0 5.0 5.0 3.0 3.0 0.6

32.1 64.9 28.2 6.0 1.0 5.6 5.5 3.2 3.3 0.6

24.2 47.6 21.3 4.4 0.8 4.0 3.7 2.4 2.6 0.5

31.0 50.0 22.0

28.0 47.0 23.0

30.0 60.0 37.0

104.0

115.0

100.0

94.0

107.0

93.0

61.0 665.0

71.0 535.0

80.0 566.0

95.0 464.0

60.0 366.0

188.0

163.0

155.0

171.0

10.0 2.0 11.0 16.0 31.0 9.0 63.0 23.0 6.0 7.0 0.0 35.0 12.0 3.0

0.0 0.0 13.0 0.0 29.0 11.9 19.0 34.0 5.0 16.0 13.0 54.0 0.0 0.0 9.52 135.22

0.0 0.0 11.0 0.0 29.0 9.6 16.0 31.0 22.0 13.0 11.0 64.0 5.0 0.0 9.85 150.49

0.0 0.0 11.0 0.0 21.0 10.8 16.0 32.0 19.0 13.0 16.0 77.0 0.0 0.0 9.33 111.49

6.44 3.30 1.35 0.47

6.66 3.39 1.39 0.44

6.30 3.46 1.24 0.49

2.40 1.09

3.34 1.15

2.24 1.02

5.84 1.57 0.64 1.62

7.31 1.65 0.61 1.25

6.47 1.16 0.86 0.99

3.56 1.22 5.50 5.89 1.56 0.64 1.70

Mean

St. dev.

71.32 11.19 3.49 1.04 0.03 0.20 1.79 2.41 0.46 0.09 2.09

17.33 3.53 1.58 0.48 0.02 0.14 0.71 1.01 0.17 0.04 0.81

34.0 52.0 42.0

28.3 45.7 27.6 4.0 0.7 3.6 3.6 2.1 2.2 0.4

8.3 15.4 9.5 2.4 0.4 2.2 2.2 1.3 1.3 0.3

97.0

99.0

95.9

41.3

82.0 521.0

82.0 747.0

50.0 344.0

75.1 417.3

32.4 185.4

216.0

154.0

174.0

202.0

169.7

50.7

13.0 2.0 10.0 21.0 26.0 11.0 81.0 35.0 9.0 19.0 9.0 49.0 14.0 3.0

9.0 2.0 10.0 14.0 30.0 8.0 47.0 19.0 6.0 9.0 0.0 50.0 11.0 3.0

10.0 3.0 12.0 18.0 30.0 9.0 63.0 23.0 7.0 12.0 0.0 50.0 12.0 3.0

13.0 3.0 11.0 43.0 30.0 10.0 88.0 46.0 19.0 29.0 28.0 100.0 17.0 2.0

8.6 2.0 10.2 15.6 26.7 8.9 51.2 26.2 8.1 11.6 5.9 51.2 10.6 2.1

4.8 1.2 2.8 10.0 7.2 2.9 27.6 14.8 5.8 6.4 7.5 22.4 6.3 1.3

2.82 0.91 5.00 6.02 1.35 0.74 1.78

3.50 1.25 5.00 6.83 1.12 0.89 1.13

3.33 1.33 4.00 7.27 1.38 0.72 1.18

3.40 1.10 3.67 5.35 1.75 0.57 1.98

3.10 1.10 4.12 6.35 1.67 0.82 1.38

1.01 0.28 1.57 2.28 1.96 0.51 0.88

136

H. B A H L B U R G

Table 4. Geochemical data for the Upper Turbidite System Sample

SiO2

SL 2

SL 9

SL 10

SL 13

SL 16

SL 17

SL 18

Mean

St. dev

A1203 Fe203 MgO MnO CaO Na20 K20 TiO2 P205 LOI Total

78.14 9.96 4.68 1.24 0.04 0.13 1.47 1.59 0.61 0.14 1.79 99.78

81.75 8.81 3.16 0.87 0.02 0.11 1.24 1.35 0.50 0.14 1.84 99.78

78.19 10.13 4.04 0.97 0.03 0.08 1.47 1.68 0.62 0.14 2.12 99.47

77.70 10.40 4.22 1.14 0.03 0.10 1.23 1.74 0.62 0.12 1.99 99.28

78.02 10.30 4.38 1.07 0.05 0.12 1.21 1.43 0.74 0.14 1.90 99.34

79.32 9.60 4.38 1.13 0.04 0.10 1.29 1.38 0.55 0.12 1.83 99.74

80.43 9.66 3.61 1.04 0.04 0.15 1.65 1.27 0.59 0.12 1.85 100.40

79.08 9.84 4.07 1.07 0.04 0.11 1.36 1.49 0.60 0.13 1.90

1.45 0.54 0.44 0.09 0.00 0.02 0.16 0.18 0.07 0.00 0.10

La Ce Nd

34.0 51.0 23.0

31.0 51.0 22.0

30.0 52.0 23.0

31.0 49.0 21.0

30.0 50.0 20.0

28.0 48.0 22.0

29.0 49.0 21.0

30.4 50.0 21.7

1.1 1.3 1.0

Rb Cs Sr Ba Hf Zr Ta Nb U Th Pb Y Sc V Cr Co Ni Cu Zn Ga Sn

78.0 16.7 47.0 350.0 5.3 221.0 0.6 12.0 3.0 11.0 8.0 25,0 9.0 101.0 39.0 12.0 17.0 15.0 43.0 13.0 3.0

83.0 9.1 660.0 1460.0 4.7 262.0 0.5 13.0 3.0 10.0 12.0 32.0 9.0 79.0 32.0 9.0 14.0 0.1 43.0 15.0 2.0

77.0 10.4 41.0 278.0 4.5 230.0 0.4 12.0 3.0 11.0 8.0 26.0 8.0 81.0 41.0 10.0 12.0 0.1 44.0 16.0 2.0

82.0 11.0 35.0 256.0 4.5 196.0 0.9 11.0 2.0 9.0 10.0 27.0 8.0 92.0 39.0 11.0 14.0 5.0 54.0 15.0 3.0

70.0 7.9 35.0 221.0 3.5 228.0 0.6 11.0 2.0 9.0 10.0 25.0 9.0 95.0 39.0 11.0 14.0 5.0 49.0 13.0 4.0

67.0 8.6 38.0 240.0 4.0 176.0 0.1 11.0 2.0 9.0 10.0 25.0 7.0 88.0 37.0 11.0 14.0 11.0 50.0 15.0 3.0

69.0 8.2 37.0 230.0 3.8 202.0 0.5 11.0 2.0 10.0 10.0 26.0 8.0 93.0 37.0 11.0 14.0 6.0 49.0 14.0 3.0

75.1 10.3 127.6 433.6 4.3 216.4 0.5 11.6 2.4 9.9 9.7 26.6 8.3 89.9 37.7 10.7 14.1 6.0 47.4 14.4 2.9

6.3 1.1 232.1 453.2 0.4 27.8 0.2 0.8 0.5 0.7 1.2 2.4 0.7 6.1 2.8 0.8 0.7 3.7 3.7 0.9 0.7

La/Sc Th/Sc Th/U SIO2/A1203 K20/Na20 Na20/K20 Rb/Sr

3.09 1.22 3.67 7.84 1.08 0.93 1.66

3.10 1.11 3.33 9,28 1.09 0.92 0.13

2.73 1.38 3.67 7.72 1.15 0.87 1.88

3.44 1.13 4.50 7.47 1.42 0.70 2.34

3.33 1.00 4.50 7.57 1.19 0.84 2.00

3.11 1.29 4.50 8.26 1.08 0.93 1.76

2.90 1.25 5.00 8.32 0.77 1.30 1.86

3.09 1.19 4.17 8.04 1.09 0.40 1.66

1.43 1.08 0.56 2.68 1.09 0.41 0.66

Comments as for Table 2. & M c L e n n a n 1985), the scatter t o w a r d s lower ratios, especially in the samples f r o m the V o l c a n o s e d i m e n t a r y Successions, is indicative o f i n p u t f r o m m a g m a t i c arc sources (Fig. 4; M c L e n n a n & T a y l o r 1991). This c o n t r a s t s with generally u p p e r crustal ratios o f > 1 in m a t u r e sands at passive margins. T h e c h o n d r i t e - n o r m a l i z e d R E E spectra have a m o d e r a t e e n r i c h m e n t o f L R E E relative to H R E E , a n d the H R E E are 7 to 20 times enriched a n d have a weakly c o n c a v e p a t t e r n (Fig. 6). T h e

negative Eu a n o m a l y is c o n s i d e r e d to be d u e to relatively a b u n d a n t plagioclase in the turbidite sandstones. Eu a n o m a l i e s ( E u / E u * ) range f r o m 0.40 to 0.67 in the V o l c a n o s e d i m e n t a r y Successions a n d are b e t w e e n 0.44 a n d 0.49 in the L o w e r T u r b i d i t e System (Tables 2, 3). A c o m p a r i s o n o f the spectra with the p o s t - A r c h a e a n average shale c o m p o s i t e ( T a y l o r & M c L e n n a n 1985) reinforces the similarity o f the O r d o v i c i a n P u n a samples to average u p p e r crustal c o m p o s i t i o n s (Fig. 6). In c o m p a r i s o n with the R E E spectra o f m o d e r n

ORDOVICIAN TURBIDITES IN THE PUNA

137

agreement with spectra of Palaeozoic active margin turbidites (Bhatia 1985; Floyd & Leveridge 1987).

Tectonic discrimination, major elements

Fig. 4. Plot of Zr/Sc v. Th/Sc (slightly modified after McLennan et al. 1993).

Fig. 5. Rb-K diagram demonstrating the close resemblance between the Puna samples and silicic and intermediate magmatic precursor rocks (after Floyd et al. 1990).

Fig. 6. REE element spectra of Ordovician Puna turbidites and the Post-Archaean average Australian shale composite (PAAS, Taylor & McLennan 1985). marine sands of variable tectonic settings, the Ordovician Puna samples compare most favourably with those of continental arcs like Java (McLennan et al. 1990). There is also a good

Several attempts have been made to base tectonic discrimination of geochemical data on major elements only (Bhatia 1983; Roser & Korsch 1986, 1988). In particular Roser & Korsch (1986, 1988) produced strong evidence to support their statistical and tectonic discrimination schemes, although there is some dependence on grain-size effects (Korsch et al. 1993). The scheme of Roser & Korsch (1988; Fig. 7) classes most of the Volcanosedimentary Succession samples as being derived from intermediate and mafic sources whereas the bulk of the remaining samples plot as being related to recycled sources. This latter result compares well with the interpretation of the framework modes presented in Fig. 2, based on the scheme of Dickinson (1985). However, the quartz content of the rocks is almost entirely of magmatic origin. The recycled orogenic signature may at least partly be influenced by the increasing but still minor input (c. 15%) from metamorphic sources (Fig. 2). Major element abundances tend to reflect the quartz content, as well as the amount of feldspar, as a function of the degree of weathering and recycling (Roser et al. 1996). The recycled sources field is commonly interpreted as representing clastic compositions typical of passive margin settings. But it seems that this particular composition reflects long periods of weathering rather than a tectonically significant passive margin signature (Roser & Korsch 1988; Roser et al. 1996). The CIA of the Puna samples indicates some major element mobility during weathering (Fig. 3), which causes the samples to plot away from the felsic and intermediate sources (Figs 5, 7) and into the recycled sources field. Samples derived from modern continental arcs are assigned correctly to mafic and intermediate sources, with one exception in field P4. The respective passive margin samples are also grouped correctly (Fig. 7). Roser & Korsch (1988) cautiously cited evidence that specific circumstances may lead to unexpected compositions of sands at active margins. Quartz-rich sediment derived from metamorphic and plutonic source terrains may be funneled into trenches and then incorporated into accretionary prisms (Velbel 1985; Bell 1987). The Ordovician Puna turbidites in turn were derived from a source of upper crustal composition represented by the evolved Ordovician

138

H. BAHLBURG

-8

-6

-4

-2

0

Discriminant

2 Function

4

6

8

1

Fig. 7. Discriminant plot of provenance and, by implication, of tectonic setting of sedimentary suites (Roser & Korsch 1988). PM, Recent deep-sea turbidites derived from and deposited at a passive margin; CA, Recent deep sea turbidites derived from and deposited at a continental arc margin, data from McLennan et al. (1990). P1, mafic sources; P2, intermediate sources; P3, felsic sources; P4, recycled sources.

arc. The lack of (meta-)sedimentary rock fragments in the Volcanosedimentary Successions, and their scarcity in the turbidite systems, argue against a volumetrically significant input from a recycled orogenic source; volcanic and plutonic detritus clearly dominates. A similar but more pronounced result is obtained from the SiO2 v. K20/Na20 discrimination of tectonic setting by Roser & Korsch

100 i ....

i ....

10

i ....

J- - , ~

9

~t..,l 1

,1

9

~

50

,

60

"-,\ 70

80

90

100

(1986; Fig. 8). The samples of the Puna turbidite systems are assigned to a passive margin setting (PCM) with most samples of the Volcanosedimentary Successions plotting in the active continental margin (ACM) and island arc fields (ARC). This discrimination is primafacie unsuccessful in the sense of tectonic setting. However, as before it can easily be explained in terms of the mobility, for example, of K and Na during weathering, of feldspar in particular. Figure 8 shows that the samples of the oldest unit, the Volcanosedimentary Successions, are mostly classed correctly as being eroded off a volcanic arc. However, as some samples plot in the passive margin field, the result is too ambiguous to be useful. The younger units are classed as recycled passive margin sandstones because they were apparently derived from source rocks longer exposed to weathering. The Recent samples are assigned correctly in broad terms. The continental arc samples plot in the active continental margin and island arc fields. Half the passive margin samples are assigned correctly with some quartz-intermediate samples in the ARC and ACM fields. The two passive margin samples in the ARC field come from the Guinea margin off western Africa and the southern part of the Falkland Islands. The Guinea sample is a passive margin sample according to it's trace and rare earth element characteristics (McLennan et al. 1990). The Falkland Islands are located on the eastern rim of the Argentinian passive margin shelf which, south of the Rio de la Plata, receives sand from the Andean arc (Potter 1994). The results obtained from the major element discrimation schemes are ambiguous at best if 'tectonic setting' is taken at face value. Many samples are erroneously classed as having been deposited at a passive margin. However, the indication of significant weathering of detritus, from the younger units in particular, explains the major element characteristics of the rocks. The evidence from the Puna samples indicates that a purely geochemical discrimination may be misleading if an estimate of the degree of weathering is not included, and that petrographical data need to be considered to balance the geochemical discrimination.

SiO2 (wt %) Fig. 8. SiO2 v. K20/Na20 discriminant plot of tectonic setting (Roser & Korsch 1986). PCM, passive continental margin; ACM, active continental margin; ARC, island arc. PM, Recent deep sea turbidites derived from and deposited at a passive margin; CA, Recent deep sea turbidites derived from and deposited at a continental arc margin, data from McLennan et al. (1990).

T e c t o n i c discrimination, trace e l e m e n t s

More promising in the evaluation of tectonic setting is the consideration of trace elements. Due to the low residence times in sea water and their characteristic behaviour during fractional crystallization, weathering, and recycling,

ORDOVICIAN TURBIDITES IN THE PUNA the HFS elements are particularly useful for provenance analysis and discrimination of tectonic setting (Bhatia & Crook 1986; McLennan 1989; McLennan et al. 1993; Roser et al. 1996). Whereas major elements reflect quartz and feldspar compositions, and weathering, immobile trace elements maintain the signature of magmatic source rocks. The ratios of La/Sc v. Ti/Zr have been shown to discriminate successfully between oceanic island arc, continental island arc, continental margin arc (active continental margin) and passive margin environments (Bhatia & Crook 1986). With six exceptions, the samples of Ordovician Puna rocks are assigned to either a continental island arc or a continental margin arc setting (Fig. 9). A similar result is obtained in the triangular plot of Sc, Zr and Th (Fig. 10), where most of the samples fall in the continental island arc field.

Discussion The interpretation of tectonic setting derived from Figs 9 and 10 is in accordance with the result inferred from petrography and the regional context, although distinction of a continental island arc form a continental margin arc setting of the Ordovician Puna turbidites cannot be made on the basis of these diagrams. However, if the data of McLennan et al. (1990)

Fig. 9. La/Sc v. Ti/Zr discriminant plot of tectonic setting (Bhatia & Crook 1986). A, oceanic island arc; B, continental island arc; C, active continental margin; D, passive margin. PM, Recent deep-sea turbidites derived from and deposited at a passive margin; CA, Recent deep sea turbidites derived from and deposited at a continental arc margin (data from McLennan et al. 1990).

139

Fig. 10. Sc-Th-Zr/10 discriminant plot of tectonic setting (Bhatia & Crook 1986). A, oceanic island arc; B, continental island arc; C, active continental margin; D, passive margin. PM, Recent deep sea turbidites derived from and deposited at a passive margin, CA, Recent deep-sea turbidites derived from and deposited at a continental arc margin (data from McLennan et al. 1990).

on modern sands are plotted in the diagrams of Bhatia & Crook (1986), doubts are cast also on the usefulness of these diagrams. In the La/Sc v. Ti/Zr diagram (Fig. 9), half the modern passive margin samples fall into the correct field D; the others plot as active continental margin or fall outside any field. The result is more promising in the ternary plot of Sc-Th-Zr/10 (Fig. 10), where only one sample plots in the arc field, with the others close to but outside the passive margin field. The continental arc samples scatter across the different arc fields in both diagrams. Bhatia & Crook (1986) showed that their discrimination scheme had a sound statistical footing regarding the samples selected. However, all samples were of Palaeozoic age and the tectonic setting of their provenance and deposition had been derived from a variety of data sets, including geochemistry. It may be more promising to approach discrimination from modern samples and settings (e.g. McLennan et al. 1990). However, it has to be emphasised that the discrimination by Bhatia & Crook (1986) is based on the concentrations of elements considered meaningful for the interpretation of tectonic setting, as they are noted for their systematic behaviour in crust-forming, magmatic processes (Taylor & McLennan 1985). In looking at the data points in Figs 9 and 10, it appears that the data base from which the

140

H. BAHLBURG

discrimination scheme was derived was perhaps too limited. Furthermore, modern passive margin sands are not as homogenous and mature as is generally expected (e.g., Valloni & Mezzardi 1984; McLennan et al. 1990; Potter 1994). Weathering and sediment dispersal paths exert a major control on the composition of these sands (Roser et al. 1996; Dickinson 1988). Thus geochemical discrimination schemes need to be used with caution if interpretation of tectonic setting is derived solely from them. A question still to be addressed is whether the Puna turbidites were derived from an island arc or from an arc resting on a continental foundation. The data presented indicate that the source of the detritus had an average upper crustal composition, and that the precursors correspond to silicic and intermediate magmatic rocks. The CIA documents significant major element changes due to weathering, which in itself is a prominent indication that major portions of the arc were exposed subaerially. More conclusive results should be obtained from a study of Nd, Sr and Pb isotope systematics that is presently under way. First results indicate that Palaeozoic sedimentary rocks including the Puna turbidites have cNdt values between - 9 . 2 and -4.9 (recalculated to the time of deposition), pointing to model ages of c. 1500 and 1900 Ma (Bahlburg et al. 1997). These data coincide with recently obtained values on Proterozoic and Palaeozoic metamorphic rocks in this region (Lucassen & Franz 1997). If confirmed by further analyses, this would suggest that the magmatic arc from which the Ordovician Puna turbidites were derived rested on relatively old continental crust, and that juvenile input from the mantle was minor (McLennan et al. 1990; Bock et al. 1996). As Neoproterozoic and Palaeozoic metamorphic rocks, granites and sedimentary rocks all have similar negative ~Ndt values and Mesoproterozoic model ages, it appears that crustal recycling played the major role, followed by the effects of weathering, in the provenance and deposition of the Ordovician Puna turbidites.

Conclusion The combination of petrographical and geochemical analyses of the Ordovician turbidites in the Puna of NW Argentina has characterized the studied rocks correctly as derived from an evolved continental margin volcanic arc of average upper crustal composition. However, the major element compositions of the Puna rocks, and in particular the KzO/Na20 ratios, mainly reflect the effects of feldspar weather-

ing. K20/Na20 ratios are low in the oldest unit (averaging 0.84), the Volcanosedimentary Successions, and increase to 1.76 and 1.09 in the Lower and Upper Turbidite Systems. This weathering effect causes the samples, which in terms of their petrography document a derivation from a volcanic source, to plot as strongly recycled sands of passive margin type in the major element discriminant diagrams (Figs 7, 8). Ratios of immobile trace elements like Zr, Ti, Sc and Th are shown to have not been altered significantly by these processes (McLennan 1989), again demonstrating their usefulness in determining the provenance of sedimentary rocks (Figs 9, 10). The determination of provenance based solely on major element data may therefore lead to erroneous results. The effect of weathering may significantly obscure the original signal. By combining the petrography of sedimentary rocks with major, trace and rare earth element geochemistry, changes due to weathering may be evaluated and used to distinguish the primary source signal from the secondary weathering influence. This study benefited from funding by Deutsche Forschungsgemeinschaft grants Gi 31/51-5 and Ba 1011/8-1. Samples AP 12 and 22 are courtesy of C. Breitkreuz. I thank N. Mortimer for his very constructive comments which led to focusing of the manuscript, and U. Jenchen and B. Bock for critical discussion. Suggestions by C. Breitkreuz and an anonymous reviewer improved the manuscript. This paper is a contribution to IGCP projects 345 'Andean lithospheric evolution' and 376 'Gondwana-Laurentia connections before Pangea'.

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H Congreso de exploracidn de Hidrocarburos, Mendoza, Actas, III, 291-307. PANKHURST, R. J., RAPELA, C. W., SAAVEDRA, J., BALDO, E., DAHLQUIST,J., PASCUA,I. 8r FANNING, C. M. 1998. The Famatinian magmatic arc in the central Sierras Pampeanas. This volume. PALMA, U. A., PARICA, P. D. & RAMOS, V. A. 1986. El granito de Archibarca: Su edad y significado tectonico, provincia de Catamarca. Revista de la Asociaci6n Geol6gica Argentina, 41, 414-419.

POTTER, P. E. 1994. Modern sands of South America. Geologisehe Rundschau, 83, 212-232. RAPELA, C. W., COIRA, B., TOSELLI,A. & SAAVEDRA,J. 1992. The Lower Paleozoic magmatism of southwestern Gondwana and the evolution of the Famatinian Orogen. International Geology Review, 34, 1081-1142. ROSER, B. P. & KORSCH, R. J. 1986. Determination of tectonic setting of sandstone mudstone suites using SiO2 content and KzO/Na20 ratio. Journal of Geology, 94, 635-650. --&--1988. Provenance signatures of sandstonemudstone suites determined using discriminant function analysis of major-element data. Chemical Geology, 67, 119-139. , MORTIMER,N., TURNBULL,I. M. & LANDIS,C. A. 1993. Geology and geochemistry of the Caples Terrane, Otago, New Zealand: compositional variations near a Permo-Triassic arc margin. In: BALLANCE, P. F. (ed.) South Pacific Sedimentary Basins. Sedimentary basins of the World, 2. Elsevier, Amsterdam, 3-19. - - - , COOPER, R. A., NATHAN, S. & TULLOCH, A. J. 1996. Reconnaissance sandstone geochemistry, provenance, and tectonic setting of the lower Paleozoic terranes of the West Coast and Nelson, New Zealand. New Zealand Journal of Geology and Geophysics, 39, 1-16. SALFITY, J. A., OMARINI, R. H., BALDIS, B. & GUTIERREZ, W. J. 1975. Consideraciones sobre la evoluci6n geol6gica del Precambrico y Paleozoico del norte Argentino. H Congreso Iberoamerica de Geologia Economica, 4, 341-361. - - - , MALANCA, S., BRANDAN, M. E., MONALDI, C. R. & MOYA, C. 1984. La Fase Guandacol en el norte de la Argentina. IX Congreso Geoldgico Argent# nos, Buenos Aires, Actas, I, 555-567. SCHWAB, K. 1973. Die Stratigraphie in der Umgebung des Salar de Cauchari (NW Argentinien). Ein Beitrag zur erdgeschichtlichen Entwicklung der Puna. Geotektonische Forschungen, 43, 1-168. TAYLOR, S. R. & MCLENNAN, S. M. 1985. The continental crust." its composition and evolution. Blackwell, Oxford. TURNER, J. C. M. 1970. The Andes of northwestern Argentina. Geologische Rundschau, 59, 1028-1063. VALLONI, R. 8z MAYNARD, J. B. 1981. Detrital modes of recent deep-sea sands and their relation to tectonic setting: a first approximation. Sedimentology, 28, 75-83. & MEZZARDI, G. 1984. Compositional suites of terrigenous deep-sea sands of the present continental margins. Sedimentology, 31, 353-364. VELBEL, M. A. 1985. Mineralogical mature sandstones in accretionary prisms. Journal of Sedimentary Petrology, 55, 685-690. YERINO, L. H. & MAYNARD, J. B. 1984. Petrography of modern sands from the Peru-Chile trench and adjacent areas. Sedimentology, 31, 83-89.

Time constraints on the Early Palaeozoic docking of the Precordillera, central Argentina VICTOR

A. R A M O S 1, R. D. D A L L M E Y E R

2 & GRACIELA

V U J O V I C H t'3

1Laboratorio de Tect6nica Andina, Universidad de Buenos Aires, (1428) Buenos Aires, Argentina 2 Department of Geology, University of Georgia, Athens, Georgia 30602, USA 3CONICET

Abstract: 4~

incremental-release ages have been determined for muscovite and hornblende concentrates prepared from basement rocks of the Sierra de Pie de Palo, in the western Sierras Pampeanas, east of the Early Palaeozoic Precordillera. The basement is mainly represented by variably metamorphosed units, including ophiolites, orthogneisses and schists. Previous U-Pb, Rb-Sr, K-Ar, and 4~ dating of magmatic and metamorphic zircons from the basement has indicated a Middle Proterozoic age. The 4~ plateau ages constrain a series of ductile deformational events that are correlated with development of structural discontinuities in adjacent foreland basins. The following deformational events are postulated: (1) initiation of collision and drowning of the Precordillera platform against Gondwana (470-460 Ma); (2) flexural extension associated with normal faulting due to tectonic loading of the Gondwana margin; and (3) development of a foreland basin (450-430 Ma). These Mid-Ordovician-Silurian events were related to collision of several exotic Early Palaeozoic terranes along the Gondwana margin. The Punta Negra foreland basin developed further west during Early to Mid-Devonian (410-380 Ma) times, and was linked to the beginning of the docking of the Chilenia terrane along the western margin of the Precordillera. Most recent workers have concluded that the Andean Precordillera represents an exotic microcontinent that collided with Gondwana during Mid-Ordovician time (see results on Penrose Conference Report, Dalziel et al. 1996). Most evidence has been obtained in the Precordillera terrane, where biostratigraphical and sedimentological studies have demonstrated that this tectonic unit was derived from Laurentia (Ramos et al. 1986; Astini et al. 1995). However, the eastern limit of the Precordillera terrane and the timing of deformation recorded in the Sierras Pampeanas have been matters of debate. The objective of this paper is to discuss time constraints for deformation of the ductile structures found in the boundary between the Precordillera and the Sierras Pampeanas. This general area is immediately west of a potential suture between Laurentia-derived terranes and the proto-margin of Gondwana. On the basis of new 4~ ages determined in the western Sierras Pampeanas, the time-span of ductile deformation is associated with docking of Early Palaeozoic terranes such as Chilenia and the Argentine Precordillera.

that developed on a Precambrian crystalline basement. These carbonate platform deposits are widely exposed between 29 ~ and 33~ within the Precordillera fold and thrust belt, as a consequence of Late Cenozoic Andean uplift (Fig. 1). Buried extensions of the carbonate platform deposits occur south of the Precordillera in isolated remnants of limestones and associated Precambrian basement. Two of the best known exposures are in the Pondn Trehue area of Mendoza and Cerro San Jorge of La Pampa province (Fig. 1). Crystalline basement blocks east of the Precordillera were uplifted during Andean deformation, and constitute the western Sierras Pampeanas geological province. Among these ranges the Sierra de Pie de Palo is important to our study, and will be discussed below. If Andean deformation is palinspastically restored, exposed basement in the western Sierras Pampeanas is seen to be in tectonic contact with the basement of Precordillera.

Regional setting

The crystalline basement of the Cambro-Ordovician carbonates is not exposed along the thin-skinned Precordillera fold and thrust belt. However, previous studies have shown that this

The Argentine Precordillera comprises carbonate shelf deposits of Cambro-Ordovician age

Precordillera basement

RAMOS, V. A., DALLMEYER,R. D. & VUJOVICH,G. 1998. Time constraints on the Early Palaeozoic docking of the Precordillera, central Argentina. In: PANKHURST,R. J. & RAPELA, C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society. London. Sr~ecial Publications. 142. 143-158.

144

V. A. RAMOS E T A L .

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Fig. 1. Regional setting of the Argentine Precordillera and western Sierras Pampeanas, with location of the study area. Location of Early Palaeozoic terranes and sutures after Ramos (1995).

basement is represented in xenoliths found in Miocene volcanic rocks of Precordillera (Leveratto 1968, 1976). Recently, Abbruzzi et al. (1993) and Kay et al. (1996) were able to date these xenoliths by U - P b zircon methods and to study their chemical and isotopic characteristics.

The xenoliths consist of amphibolites, and biotite-bearing acidic gneisses associated with mafic pyroxene granulites and pyroxene-garnet granulite gneisses. The chemistry of these rocks indicates a complex oceanic island arc or backarc protolith; it includes a mafic igneous suite of

EARLY PALAEOZOIC PRECORDILLERA DOCKING relatively unfractionated mantle-derived arc magmas (Kay et al. 1996; Vujovich & Kay this volume). The Precordillera xenoliths contain the least radiogenic Pb of Precambrian to Recent rocks in South America (Tosdal et al. 1994; Kay et al. 1996; Tosdal 1996). The Pb signatures are also strikingly similar to those of the North American Grenville Province, and were considered by these authors as diagnostic of Eastern

145

Laurentian derivation. Kay et al. (1996) report U-Pb zircon ages: prismatic zircons gave 1102 + 6 Ma for a mafic xenolith and 1118 + 54 Ma for an acidic one, whereas the maximum age of metamorphism based on clear rounded zircons was estimated as c. 1083 Ma. These ages are similar to those reported for basement rocks of the Sierra de Pie de Palo by McDonough et al. (1993).

LEGEND

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Fig. 2. Simplified geological map of the Pie de Palo region (after Ramos & Vujovich 1995). See inset in Fig. 1.

146

v. A. RAMOS ET AL.

Pie de Palo basement

Basement o f the Cuyo Basin

The crystalline rocks of Sierra de Pie de Palo, Cerro Barboza and Cerro Valdivia (Fig. 2) form the largest continuous exposures of the basement adjacent to Precordillera. Different authors have described this basement (see recent synthesis of Vujovich & Kay 1996). It was divided into two metamorphic units, the Pie de Palo complex and the Caucete Group, by Ramos & Vujovich (1995). The Pie de Palo complex consists in part of an ophiolitic assemblage, represented by peridotites, meta-gabbros, greenschists, talc, serpentinites and amphibolites, associated with mica and graphite schists. East of the ophiolites, the central and eastern parts of the Sierra de Pie de Palo are composed of feldspar-biotite-garnet schists, biotite-muscovite-garnet-plagioclase schists and gneisses, amphibolites and marbles. The Caucete Group (sensu Ramos & Vujovich 1995) consists of low-grade metamorphic rocks; these were carbonate and clastic shelf sequences, now represented by marbles, quartzites and quartzose schists. The Caucete Group also contains limestones, penetratively deformed along the Angaco belt on the western side of Sierra de Pie de Palo. The contact between the Caucete Group and the Pie de Palo complex is tectonic (Fig. 2), and it has been mapped along most of the range (Vujovich & Ramos 1994). The metamorphic rocks of the Pie de Palo basement grade from greenschists to amphibolite in western sectors, whereas in the east the metamorphic grade increases up to high amphibolite, locally granulite, facies (Dalla Salda & Varela 1984; Ramos & Vujovich 1995). The mafic and ultramafic rocks are of island-arc and back-arc oceanic affinities, that developed into more sialic signatures eastward, where the arc might have developed in attenuated continental crust (Vujovich & Kay 1996). U-Pb ages of abraded zircons from magmatic and metamorphic rocks of the Pie de Palo complex have indicated Middle Proterozoic ages (McDonough et al. 1993). Orthogneiss zircons have an upper intercept of 1060-+-20 Ma, which was interpreted as the age of crystallization (McDonough et al. 1993); recent recalculation suggests an older intercept of 1105• Fractions from amphibolites and meta-rhyolites gave nearly concordant data with 2~176 ages of 1060-1080 Ma. The amphibolite zircons are platelets, typical of metamorphic zircons from mafic rocks. The coincidence of 2~176 ages for these samples suggests that they record a medium- to high-grade metamorphic event at that time (McDonough et al. 1993).

Buried extensions of the Precordillera carbonate shelf are covered by Triassic continental clastic deposits of the fault-bounded Cuyo basin. This basin has two depth-centres, the Cacheuta and the General Alvear sub-basins (Fig. 1). Drilling associated with oil exploration has revealed that beneath the Triassic deposits of the Cacheuta sub-basin in the northern Cuyo area (Fig. 1), there are widely developed Cambro-Ordovician carbonates and metamorphic Precambrian rocks. The distribution of these basement rocks is complex, possibly reflecting either Famatinian (Early Palaeozoic) deformation, or Triassic normal faulting. South of the General Alvear subbasin, basement has also been intersected at shallow depths by drilling. Four contrasting lithological assemblages have been penetrated: (1) Cambro-Ordovician carbonate deposits typical of the Precordillera shelf; (2) Early Palaeozoic clastic deposits, equivalent to Punta Negra foreland basin (Astini et al. 1996); (3) low-grade metamorphic rocks of the Angaco belt (Vujovich & Ramos 1994); and (4) crystalline basement, mainly correlated with the Pie de Palo complex (Pucci 1976). These basement rocks comprise an apparently coherent tectonic element that extends south of the Precordillera and Sierra de Pie de Palo, and is bounded to the west by mafic and ultramafic rocks.

Pon6n Trehue basement

These isolated exposures of basement rocks have been assigned to Cerro La Ventana Formation by Criado Roque (1972), and described by Ardolino (1973) and Nfifiez (1979). They consist of amphibolites, quartzose mica schists and gneisses intruded by tonalite, diorite, and other granitoids. Their metamorphic facies correspond to albite-epidote amphibolite and they have undergone conspicuous cataclastic deformation. The basement rocks are overlain by carbonate platform deposits with fossils that correlate with the Precordillera carbonate shelf (Bordonaro et al. 1996; Heredia 1996).

Cerro San Jorge area

The limestones of Cerro San Jorge are penetratively deformed. Although their basement is not exposed, there are several crystalline rocks exposures between Ponon Trehue and San Jorge (Wichmann 1928; Holmberg 1973). This

EARLY PALAEOZOIC PRECORDILLERA DOCKING poorly exposed basement, which has been confirmed by drilling, consists of granitoids, biotite schists and gneisses. Garnet-bearing schists were reported in several boreholes and yielded wholerock K - A r ages of 605 Ma (Pucci 1976), regarded as a minimum age. Linares et al. (1980) reported K - A r ages of 770 + 40 Ma for dioritic rocks cropping out in the vicinity. Several authors have correlated the San Jorge limestones either with the carbonate shelf of the Precordillera (Wichmann 1928) or with the Caucete Group (Astini et al. 1995).

The Cuyania composite terrane Based on petrographical, geochemical, and isotopic characteristics, the basement of Precordillera and that of the Pie de Palo have been considered to represent a single block of Laurentian derivation (Abbruzzi et al. 1993; Tosdal et al. 1994; Kay et al. 1996). Fossiliferous carbonate shelf deposits overlying the basement of Ponon Trehue have a trilobite fauna that correlates with those of the Precordillera (Baldis & Blasco 1973; Bordonaro et al. 1996). Furthermore, the subsurface remnants of carbonate and basement rocks of the Cuyo basin strongly support the continuity of the basement between the Precordillera and Ponon Trehue. If correlation between the San Jorge area and the Sierra de Pie de Palo is accepted, as proposed by Astini et al. (1995), it suggests that a large area of basement was involved in the collision of the Precordillera. This Early Palaeozoic composite terrane, termed Cuyania by Ramos (1995), exceeds 1000 km in length. It is considered to be a composite terrane because an ophiolitic belt of Middle Proterozoic age separates the Precordillera and the Pie de Palo areas (Fig. 2). As noted above, the metamorphic basement of the Cuyania composite terrane has been dated in the Precordillera and Sierra de Pie de Palo regions by Kay et al. (1996) and McDonough et al. (1993), respectively. Re-assessment of the U-Pb zircon data indicates that a magmatic component of the Sierra de Pie de Palo basement crystallized during Grenville orogenesis at c. 1100 Ma, followed by a medium to high-grade metamorphism, possibly as much as c. 30 to 35 Ma later. These age data compare to syntectonic magmatism from 1160 to about 1060 Ma followed by high-grade metamorphism at about 1040Ma in several Grenville areas, such as in southern Ontario (McEachern et al. 1990). Therefore, it seems reasonable to assume a

147

Grenville age for the main igneous and metamorphic events in the Cuyania terrane. These basement rocks were also affected by a penetrative ductile deformation, probably during accretion of the Precordillera. The age of this last episode has been poorly constrained until now due to the lack of suitable ages for the deformation.

Pre-Andean structure The ductile deformation that affected the basement is well exposed in the Sierra de Pie de Palo, where it has been studied by different workers (Dalla Salda & Varela 1984; Ramos & Vujovich 1995). One of the most outstanding features is the overthrust that juxtaposes rocks of the Caucete Group and the Pie de Palo complex, known as the Las Pirquitas thrust (Ramos & Vujovich 1995). This tectonic contact has been observed at several locations in the western foothills of Sierra de Pie de Palo (see Fig. 2). It carries basic and ultrabasic metamorphic rocks of the ophiolitic complex westwards over the platform metasediments. The thrust front shows a set of imbricated splays, with dips varying from horizontal up to 20-30~ In the vicinity of the fault plane, on both the footwall and the hangingwall, strong ductile shearing is observed. The fault planes have marked stretching lineations that vary in direction from N80~ to N110~ Kinematic indicators show a dominant transport to the northwest, with local components to the north and southwest. A series of ductile deformed west-vergent thrust slices occurs east of the main overthrust. These thrusts are developed in mylonitic gneisses of tonalitic to granitic composition, with intercalated lenses of amphibolite and quartzK-feldspar-muscovite schists derived from rhyolites. These rocks were thrust westward onto a package of mixed orthogneisses and paragneisses, also with meta-rhyolite and amphibolite lenses. All these rocks override the Caucete Group rocks along the western side of the range. These structures are all intersected by a preAndean erosional peneplain, which was not modified by the thrusting. Studies of this peneplain in different sectors of the Sierras Pampeanas suggest a minimum Late Palaeozoic age for the beginning of its development (Jordan et al. 1989).

Analytical methods The Sierras Pampeanas mineral concentrates were analysed using incremental-release, 4~ analysis (Figs 3 and 4, Tables 1-3). The techniques used

148

V. A. R A M O S E T AL.

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% 39Ar RELEASED

Fig. 4. 4~ apparent age spectra of muscovite concentrates from the Sierras Pampeanas, Argentina. Data plotted as in Fig. 3. generally followed those described by Dallmeyer & GilIbarguchi (1990). Variations in flux of neutrons along the length of the irradiation assembly were monitored with several mineral standards, including MMhb-1 (Sampson & Alexander 1987). Intra-laboratory uncertainties have been calculated by statistical propagation of uncertainties associated with measurements of each isotopic ratio (at two standard deviations of the mean) through the age equation. Inter-laboratory uncertainties are about +1.25-1.5% of the quoted age. A 'plateau' is considered defined if the ages recorded by four or more contiguous gas fractions (with similar apparent K/Ca ratios), each representing >4% of the total 39Ar evolved (and together constituting >50% of the total 39Ar evolved), are mutually similar within + 1% intra-laboratory uncertainty. Analyses of

the MMhb-1 monitor show that apparent K/Ca ratios may be calculated through the relationship 0.518(+0.005) x (37Ar/39Ar)..... cted. Plateau portions of the analyses were plotted on 36Ar/4~ v. 39Ar/4~ isotope correlation diagrams. Regression techniques followed methods described by York (1969), the mean square of weighted deviates (MSWD) being used to evaluate the isotope correlations.

Results

Hornblende The four h o r n b l e n d e concentrates display internally discordant 4~ age spectra of variable

V. A. R A M O S E T AL.

150

Table 1. 4~ analytical data for incremental-heating experiments on hornblende concentrates from the Sierras Pampeanas, Argentina Realease temp (~

(4~

Sample 104: J =-0.009661 610 117.19 710 65.15 735 42.97 760 42.74 785 35.36 810 32.41 835 31.79 860 31.86 885 29.34 905 29.26 930 30.76 955 31.72 990 30.77 1025 31.55 1060 31.73 Fusion 37.09 Total 35.71 Total without 610-905~ 1060~ Sample 106: J = 0.009480 610 99.43 710 76.24 735 78.60 760 63.28 785 53.56 810 51.39 835 49.64 860 55.80 890 54.11 920 49.12 Fusion 34.71 Total 63.44 Sample 108: J = 0.009755 610 38.27 710 39.54 735 44.08 760 42.67 785 32.88 810 30.75 835 28.69 860 29.85 885 33.58 910 33.29 935 41.72 965 34.30 Fusion 36.71 Total 35,09 Total without 610-760~ 965~ Sample 109." J = 0.009665 610 42.08 710 37.68 735 33.97 760 29.79 785 31.94

(36Ar/39Ar)*

(37Ar/39Ar)C

39Ar % of total

%4~ non-atmos.t

36Area %

Apparent Age (ma)~c

0.21087 0.09918 0.02458 0.02476 0.01523 0.01417 0.01884 0.01752 0.01772 0.01320 0.00588 0.00945 0.00538 0.00854 0.00407 0.02021 0.01965

3.990 11.674 12.519 12.436 12.297 10.657 14.211 12.165 11.444 11.804 12.815 13.541 13.582 13.897 12.905 17.167 12.951

3.37 2.77 1.68 1.69 1.10 1.29 1.68 6.38 4.78 5.22 8.52 13.95 13.23 22,56 7.35 4.43 100.00 58.27

47.10 56.45 85.42 85.20 90.05 89.71 86.06 86.80 85.27 89.89 97.68 94.61 98.37 95.52 99.46 87.60 91.20

0,51 3.20 13.85 13.66 21.97 20.46 20.52 18.89 17.57 24.32 59.32 39.00 68.72 44.24 86.17 23.10 41.66

772.54-2.0 551.9 4- 0.7 551.2 • 0.9 547.5 4- 0.9 487.l 4- 1.6 449.3 4- 1.3 426.5 4- 0.9 430.04-0.4 393.1 4- 1.1 411.1 4- 1.3 463.1 4-0.5 462.8 4- 0.2 466.3 4- 0.4 464.6 4- 0.2 483.44- 1.1 497.24-2.5 473.9 4- 0.7 464.3 4- 0.3

0.09232 0.09922 0.11509 0.06375 0.03315 0.06626 0.06918 0.08709 0.08257 0.06628 0.01382 0.07389

2.896 7.2t3 7.316 10.186 12.092 12.401 10.845 11.567 8.521 8.922 10.407 9.216

9.49 13.23 12.38 11.83 12.75 9.55 8.95 6.24 4.40 5.44 5.75 100.00

72.79 62.29 57.47 71.52 83.51 63.83 60.56 55.54 56.17 61.57 90.63 67.38

0.85 1.98 1.73 4.35 9.92 5.09 4.26 3.61 2.81 3.66 20.49 4.93

943.8 4- 0.6 673.1 • 645.5 4- 0.8 647.6 4- 0.7 641.9 4- 0.6 491.8 • 0.7 455.1 • 0.5 467.7 4- 1.0 459.04-2.1 457.0 4- 1.0 473.6 4- 0.2 606.1 4- 0.7

0.05849 0.06233 0.06322 0.06849 0.03773 0.03032 0.02453 0.02787 0.04196 0.03964 0.06848 0.03437 0.04126 0.04447

2.104 5.517 10.133 11.466 11.450 13.395 13.586 13.846 15.581 14.638 14.987 10.387 11.534 11.278

10.35 9.89 6.82 4.49 5.93 8.80 10.21 16.47 7.91 4.42 7.35 4.07 3.27 100.00 61.09

55.26 54.53 59.45 54.72 68.87 74.35 78.52 76.12 66.79 68.33 54.37 72.81 69.29 66.35

0.98 2.41 4,36 4.55 8.25 12.02 15.06 13.51 10.10 10,06 5,95 8.22 7,60 8.42

338.8 4- 0.4 345.5 + 0.3 413.0+0.2 372.6 4- 0.7 362.5 4- 0.4 366.04-0.4 361.2 4- 0.3 364.1 +0.3 360.1 4- 0.3 364.6 4- 1.1 363.7 4- 0.6 395.64-0.6 402.5 4- 0.7 365.5 4- 0.4 363.2 • 0.4

0.05652 0.03181 0.01258 0.00275 0.01547

1.424 1.104 1.484 2.369 3.910

4.66 9.59 8.44 3.30 3.22

60.57 75.27 89.39 97.90 86.65

0.69 0.94 3.21 23.46 6.87

397.6 4- 0.8 437.34-0.2 464.7 4- 0.2 448.6 4- 0.4 428.64-0.7

EARLY PALAEOZOIC PRECORDILLERA

DOCKING

151

Table 1. (continued) Realease temp (~

(4~

(36Ar/39Ar)*

Sample 109: J = 0.009665 (continued) 810 30.33 0.01541 835 30.02 0.01189 855 30.82 0.01405 875 29.70 0.00980 895 29.37 0.00858 920 29.51 0.00992 950 29.99 0.01251 985 29.25 0.00400 Fusion 29.57 0.00365 Total 31.33 0.01462 Total without 610-785~ fusion Sample l l h J =0.009673 610 66.26 710 76.39 735 58.33 760 36.63 810 72.54 835 32.19 855 67.68 875 64.20 895 48.32 920 99.45 940 83.48 970 79.93 1000 79.99 1040 82.68 Fusion 74.78 Total 73.66

0.15648 0.18220 0.08941 0.02356 0.16302 0.01645 0.14013 0.10904 0.04741 0.13648 0.06298 0.03640 0.01946 0.01968 0.04438 0.06117

(37Ar/39Ar)C

39Ar % of total

%40At non-atmos.t

36Arca %

5.544 7.247 7.797 8.484 8.646 9.186 9.724 7.834 9.142 6.433

3.70 6.78 9.70 10.74 11.14 11.01 8.78 5.98 2.95 100.00 69.35

86.43 90.21 88.54 92.53 93.71 92.55 90.26 97.79 98.82 88.77

9.78 16.57 15.09 23.55 27.40 25.19 21.14 53.25 68.13 19.75

428.6 • 0.5 421.2+0.3 424.2 + 0.3 426.9 + 0.2 427.6 + 0.2 424.9+0.2 421.6 + O.4 420.9 + 0.6 451.1 + 1.3 426.9 + 0.4 424.7 + 0.3

1.034 6.736 3.289 3.516 6.106 2.365 23.475 27.138 23.644 60.430 33.340 21.098 16.159 14.284 13.936 15.725

6.46 6.84 5.51 2.74 2.93 3.78 1.24 2.31 4.45 3.04 5.53 11.35 19.32 20.70 3.84 100.00

30.33 30.22 55.14 81.75 34.26 85.47 41.59 53.20 74.93 64.32 80.91 88.66 94.43 94.35 83.95 76.08

0.18

320.6• 364.9 + 2.0 489.4 • 2.0 459.7 + 2.8 389.9 -4-6.7 426.1 + 1.8 44O. 1-4- 3.9 522.7 • 9.1 548.6 + 8.9 895.9 + 10.2 922.1 + 5.4 951.5+ 1.1 997.0 + 1.8 1021.1 + 1.9 861.9 • 8.4 777.8 + 3.1

Apparent Age (ma):~

1.01 1.00

4.06 1.02

3.91 4.56 6.77 13.57 12.04 14.40 15.76 22.59 19.74 8.54 12.97

* Measured. c Corrected for post-irradiation decay of 37Ar (35.1 day 89 t [4~ .... )(295-5)]/4~ . Two sigma, intralaboratory errors. Table 2. 36Ar/4~ v. 39Ar/4~ isotope correlations for incremental-heating experiments on hornblende concentrates from the Sierras Pampeanas, Argentina Sample

104 108 109

4~

Isotope correlation age (Ma)*

interceptt

436.1+0.9 360.4 -t- 0.8 423.7 + 0.9

292.7-t-7.8 294.8 • 3.7 296.3 + 4.2

MSWD

1.37 1.96 1.08

Increments included (~

% of total 3 9 A r

4~

Calculated plateau age (Ma)

930 1025 785-935 810-985

58.27 61.09 65.35

404.3+0.3 363.3 + 0.4 424.7 4- 0.3

Calculated using the inverse abscissa intercept (4~ * Inverse ordinate intercept. t Table 1.

ratio) in the age equation.

complexity. All spectra are m a r k e d by considerable v a r i a t i o n s in a p p a r e n t ages r e c o r d e d by gas fractions evolved at low a n d very high experim e n t a l t e m p e r a t u r e s . These are m a t c h e d by fluctuations o f K / C a ratios t h a t suggest t h a t A r was evolved f r o m c o m p o s i t i o n a U y distinct,

relatively n o n - r e t e n t i v e phases. These c o u l d be r e p r e s e n t e d by: (1) very m i n o r , optically u n d e t e c t a b l e m i n e r a l o g i c a l c o n t a m i n a n t s in the h o r n b l e n d e concentrates; (2) p e t r o g r a p h i cally u n r e s o l v e d e x s o l u t i o n or c o m p o s i t i o n a l z o n a t i o n within c o n s t i t u e n t h o r n b l e n d e grains;

152

V. A. R A M O S

E T AL.

Table 3. 4~Ar/39 Ar analytical data for incremental-heating experiments on muscovites concentrates from the Sierras Pampeanas, Argentina Realease temp (~

(4~

Sample 101." J = 0.009605 510 42.38 575 25.66 600 25.90 610 25.73 640 25.80 670 25.59 690 26.00 710 25.93 730 25.81 760 25.77 790 26.03 825 25.91 860 26.02 900 26.05 940 25.87 980 26.08 Fusion 24.82 Total 26.06 Total without 510-575~ and fusion Sample 102: J = 0.010181 510 32.30 550 28.39 580 27.33 600 27.34 620 27.24 640 27.16 660 27.19 680 27.23 705 27.26 730 27.25 760 27.26 790 26.64 820 26.24 Fusion 26.09 Total 27.49 Total without 510-550~ 790~ Sample 103: J = 0.008943 510 36.42 550 32.18 580 30.27 600 29.92 620 30.04 640 29.56 660 29.57 685 29.45 710 29.50 740 29.70 780 29.51 820 29.31 Fusion 29.62 Total 29.96 Total without 510-580~ Sample 105: J = 0.009125 510 25.50 550 27.64

(36Ar/39Ar)*

(37Ar/39Ar)C

39mr % of total

%4~ non-atrnos.t

36Arca %

Apparent Age (ma);

0.05253 0.00118 0.00796 0.00095 0.00011 0.00093 0.00226 0.00095 0.00045 0.00154 0.00174 0.00162 0.00146 0.00121 0.00197 0.00108 0.00195 0.00199

0.023 0.020 0.008 0.019 0.037 0.010 0.008 0.012 0.024 0.014 0.009 0.016 0.027 0.009 0.023 0.018 0.143 0.018

0.83 1.95 2.85 1.71 2.20 3.09 2.74 9.68 3.59 5.55 10.37 7.42 9.89 13.47 13.48 10.02 1.17 100.00 93.20

63.36 98.62 90.90 98.90 99.86 98.91 97.41 98.90 99.47 98.22 98.01 98.14 98.33 98.61 97.73 98.76 97.70 97.90

0.01 0.45 0.03 0.55 9.04 0.30 0.09 0.34 1.45 0.25 0.14 0.26 0.51 0.21 0.32 0.44 2.00 0.55

414.0 • 0.8 392.5 • 0.1 367.7 • 0.1 394.5 • 0.2 398.8 • 0.4 392.6 • 0.2 392.8 + 0.2 397.2 • 0.2 397.6 • 0.1 392.5 • 0.3 395.5 • 0.2 394.2 + 0.3 396.5 • 0.2 397.9 • 0.2 397.1 • 398.7 4- 0.2 377.7 ~=0.9 394.7 • 0.2 395.7 • 0.2

0.02536 0.00327 0.00308 0.00233 0.00161 0.00190 0.00174 0.00211 0.00239 0.00208 0.00216 0.00458 0.00313 0.00712 0.00321

0.027 0.026 0.017 0.022 0.014 0.015 0.023 0.023 0.008 0.016 0.012 0.023 0.013 0.005 0.017

5.64 3.82 8.32 5.16 4.52 6.61 9.50 8.27 9.60 13.88 15.43 3.06 3.29 2.91 100.00 81.29

85.94 96.58 96.66 97.47 98.23 97.92 98.09 97.70 97.39 97.72 97.64 94.91 96.46 91.91 96.64

0.05 0.21 0.15 0.26 0.24 0.21 0.35 0.30 0.09 0.21 0.15 0.14 0.11 0.02 0.19

448.9 4- 0.3 444.1 + 0.2 429.7 + 0.2 433.0 4- 0.3 434.6 • 0.2 432.3 • 433.3 + 0.3 432.3 9 0.2 431.6 • 0.2 432.8 zt: 0.3 432.6 • 0.2 413.3 • 0.1 413.6• 394.0 • 0.4 431.4• 432.4 • 0.2

0.02166 0.00763 0.00271 0.00315 0.00394 0.00246 0.00230 0.00267 0.00274 0.00293 0.00260 0.00246 0.00306 0.00364

0.017 0.026 0.013 0.017 0.018 0.027 0.024 0.007 0.010 0.010 0.008 0.011 0.008 0.014

3.97 3.24 4.99 4.13 4.58 8.68 8.10 15.92 11.38 12.29 11.71 9.03 1.98 100.00 87.79

82.41 92.98 97.33 96.87 96.11 97.53 97.69 97.30 97.24 97.07 97.38 97.50 96.93 96.53

0.02 0.09 0.13 0.15 0.12 0.30 0.28 0.07 0.10 0.09 0.09 0.12 0.07 0.13

428.8 + 0.3 427.7 + 0.2 421.8 4- 0.4 415.7• 414.2 i 0.2 413.7 • 0.2 414.4 ~ 0.3 411.4 • 0.3 411.9:t=0.4 413.7 :t: 0.2 412.4 • 0.3 410.5 :k 0.1 412.1 J: 0.2 414.3 ~ 0.3 412.7 & 0.3

0.05573 0.00400

0.060 0.016

1.06 2.63

35.42 95.71

0.03 0.11

142.9 :t: 0.3 390.1 :t: 0.2

EARLY PALAEOZOIC

PRECORDILLERA

DOCKING

153

Table 3. (continued) (37Ar/39Ar)c

39At % of total

%40Ar non-atmos.t

36Arca %

Apparent Age (ma);

Sample 105: J = 0.009125 (continued) 600 28.57 0.00316 650 27.77 0.00398 680 27.44 0.00265 710 27.46 0.00188 740 27.52 0.00222 770 27,39 0.00157 800 27,38 0.00151 830 27.30 0.00128 860 27.27 0.00149 900 27.23 0.00171 940 27.23 0.00131 980 26.78 0.00054 1020 27.05 0.00125 1060 27.28 0.00141 Fusion 27.24 0.00519 Total 27.32 0.00237 Total without 510-550~

0.011 0.019 0.016 0.013 0.006 0.028 0.007 0.006 0.015 0.010 0.015 0.007 0.010 0.012 0.033 0.012

6.21 3.34 3.39 3.36 14,46 2.77 4.67 3.79 2.93 6.05 7.47 13.19 13.39 9.51 1.77 100.00 88.33

96.71 95.75 97.12 97.96 97.59 98.29 98.35 98.59 98.37 98.12 98.56 99.39 98.62 98.45 94.36 97.38

0.09 0.13 0.16 0.18 0.07 0.48 0.13 0.14 0.27 0.15 0.31 0.33 0.21 0.23 0.17 0.20

405.6 • 0.2 391.8 4- 0.2 392.7 4- 0.1 396.0 + 0.2 395.5+0.2 396.2 4- 0.2 396.4 • 0.1 396.2 4- 0.2 395.1 +0.4 393.6 + 0.2 395.2 + 0.2 392.2 4- 0.1 393.1 +0.3 395.5+0.4 380.1 + 0.2 392.0 • 0.2 394.3 4-0.2

Sample 107: J = 0.009450 510 30.25 550 27.77 580 27.08 600 26.78 620 26.78 640 26.80 660 26.92 685 26.95 710 27.00 740 27.08 780 27.14 820 27.22 Fusion 27.33 Total 27.34 Total without 510-550~

0.00808 0.00101 0.00095 0.00041 0.00092 0.00126 0.00036 0.00139 0.00097 0.00146 0.00106 0.00099 0.00161 0.00170

0.035 0.033 0.028 0.050 0.061 0.060 0.054 0.054 0.035 0.059 0.045 0.036 0.076 0.048

8.91 8.55 7.42 5.85 6.62 10.47 6.28 9.38 8.95 14.94 4.69 5.13 2.81 100.00 82.54

92.09 98.91 98.95 99.54 98.98 98.61 99.60 98.47 98.92 98.40 98.84 98.91 98.26 98.22

0.12 0.89 0.81 3.26 1.79 1.29 4.15 1.05 0.98 1.10 1.17 0.98 1.29 1.34

421.6+0.3 416.3+0.1 407,2 • 0.3 405.3 • 0.4 403.2 + 0.2 402.2 + 0.3 407.4 ~: 0.3 403.7 • 0.2 406.0 • 0.2 4O5.2 + 0.3 407.6 + 0.2 408.9 ~ 0.2 408.0 + 0.2 407.8 + 0.3 405.4 + 0.3

0.07381 0.00448 0.00305 0.00203 0.00200 0.00253 0.00150 0.00067 0.00080 0.00006 0.00053 0.00060 0.00121 0.00127 0.00432 0.00308 0.00236

0.208 0.224 0.276 0.023 0.010 0.010 0.008 0.008 0.011 0.012 0.007 0.007 0.012 0.014 0.210 0.033 0.048

1,15 4.23 8.04 4.97 5.88 5.39 5.24 7.00 8.12 12.36 11.19 8.23 8.75 3.86 2.04 3.55 100,00 61.54

54.51 95.19 96.83 97.81 97.85 97.29 98.40 99.28 99.15 99.35 99.44 99.36 98,73 98.68 95.69 97.05 97.88

0.08 1.36 2.46 0.31 0.14 0.11 0.14 0.33 0.36 0.55 0.37 0.33 0.27 0.29 1.32 0.29 0.56

396.4 + 0.6 394.5 -t- 0.2 407.8 + 0.2 407.4 + 0.1 408.9 + 0.2 408.7 + 0.2 414.0+t-0.2 421.3 + 0.1 424.8 + 0.1 425.0 + 0.3 427.0 + 0.4 427.2 + 0.2 427.2 + 0.2 427.5 + 0.2 427.3 + 0.3 447.9 • 0.3 419.6+0.2 425.0 + 0.2

Realease temp (~

(4~

Sample 110: J =0.009405 510 47.92 550 27.30 600 27.85 650 27.55 680 27.64 710 27.79 740 27.87 770 28.17 800 28.47 830 28.43 860 28.55 890 28.59 930 28.47 970 28.81 1010 29.38 Fusion 30.87 Total 28.54 Total without 510-740~ and fusion

(36Ar/39Ar)*

* Measured. c Corrected for post-irradiation decay of 37Ar (35.1 day 89 t [40Artot-36Arat. . . . )(295-5)]/4~ 9 Two sigma, intralaboratory errors.

154

V. A. RAMOS E T AL.

(3) minor chloritic replacement of hornblende; and/or (4) intra-crystalline inclusions. Except for sample 111, the intermediate-temperature gas fractions measured from the hornblende concentrates display little intra-sample variation in apparent K/Ca ratios suggesting Ar release from populations of compositionally uniform intracrystalline sites. The intermediate-temperature fractions evolved in the analyses of hornblende concentrates 104, 108, and 109 display little intra-sample age variation, and define plateaux corresponding to 464.3 +0.3, 363.2:1:0.4 and 424.7 + 0.3 Ma. The plateau analytical data yield well-defined 36Ar/4~ v. 39Ar/4~ plateau isotope correlations (MSWD < 2.0), with inverse ordinate intercepts (4~ similar to that of the 4~ ratio in the present-day atmosphere (Table 2). This suggests that there is no significant intra-crystalline contamination with extraneous ('excess') argon components. Using the inverse abscissa intercepts (4~ in the age equation yield plateau isotope-correlation ages of 463.1 :t:0.9 (104), 360.4+0.8 (108) and 423.7+0.9Ma (109). Because calculation of isotope-correlation ages does not require assumption of present-day 4~ ratios, they are considered more reliable than ages calculated directly fi'om the analytical data. They are interpreted as dating the last cooling through temperatures required for intra-crystalline retention of argon in the constituent hornblende grains. Harrison (1981) suggested that temperatures of c. 500 + 25~ are appropriate for argon retention within most hornblende compositions in the range of cooling rates likely to characterize most geological settings. The hornblende concentrates prepared from samples 106 and 111 display complex, internally discordant apparent age spectra which correspond to total-gas ages of 606.1 + 0 . 7 M a (106) and 777.8• (111). Plateaux are not defined, and no meaningful isotope correlations are defined by any combination of the data. Therefore no significance is ascribed to the totalgas ages. The high temperature increments evolved from sample 111 record systematically increasing ages which approach 1050 Ma, but a plateau is not statistically defined. The possible geological significance of the 1050Ma age is discussed later.

Muscovite

The six muscovite concentrates display variably discordant 4~ age spectra. Apparent K/Ca ratios are very high, with considerable

uncertainties. Consequently they are not shown with the age spectra. The apparent age spectra from samples 101, 102, 103, 105 and 107 are characterized by slight intra-sample age variations at low temperatures, but display well-defined intermediateand high-temperature plateaux. These range between 432 + 0.2Ma (102) and 394.3 4- 0.2 Ma (105), and are all interpreted as dating the last cooling through temperatures required for intracrystalline retention of argon within constituent muscovite grains. Closure temperatures of c. 400 + 25~ have been considered appropriate for muscovite by Cliff (1985) and Blackenburg et al. (1989). The muscovite concentrated prepared from sample 110 is characterized by more extensive low- and initial-intermediate temperature age discordance, but the 800-1010~ increments record similar apparent ages which correspond to a plateau age of 4254-0.2Ma, which is considered a geologically significant cooling age.

Geological interpretation The results obtained from muscovite concentrates from a variety of rock types, in different tectonic settings of the Sierra de Pie de Palo, are thought to indicate a cooling age associated with the last episode of ductile deformation and uplift. Sample 102 is a biotite-muscovite-garnet schist, representative of the central part of Sierra de Pie de Palo (Fig. 2). These rocks show protomylonitic textures, developed along fault zones. Porphyroclasts of quartz, feldspar and garnet are immersed in a mylonitic matrix composed of dynamic recrystallized quartz layers that developed a weakly oblique foliation. The phyllosilicate-rich layers show abundant biotite intergrowth with muscovite crystals, together with quartz and feldspar. Intra-crystalline deformation produced undulose extinction and bending. The cooling of this last episode of deformation may be represented by the 432 + 0.2 Ma age. Further south, quartz-plagioclase-muscovite gneisses, such as sample 105, are also mylonitized. They are characterized by a matrix of recrystallized quartz with small and idiomorphic subparallel muscovite crystals. The muscovite concentrate age of 394.3 + 0 . 2 M a is probably associated with this ductile deformation. Some small bodies of tonalite in southeastern sector of Pie de Palo (Fig. 2), are represented by protomylonitic rocks. Sample 107 comes from one of these in the Santa Rosa area. Abundant

EARLY PALAEOZOIC PRECORDILLERA DOCKING plagioclase porphyroclasts occur in a matrix with mylonitic foliation. The matrix is composed of quartz-feldspar recrystallized aggregates and thin layers of recrystallized muscovite, with undulose extinction and bending. The muscovite age of 405.4 + 0.3 Ma may represent the protomylonite cooling age. Strongly deformed leucogranites, such as sample 103, are characterized by recrystallized quartz aggregates, alternating with bands of idiomorphic muscovite sub-grains. The deformation of these granites may have a cooling age of 412.7 -t- 0.3 Ma (Fig. 4). Thus, Ar-Ar muscovite ages of the Pie de Palo complex, located in the hanging wall of the Las Pirquitas thrust, independently of the lithology, vary from 432 to 394 Ma. The quartzites of the Caucete Group, exposed in the footwall of the Pirquitas thrust, are mainly composed of quartz and muscovite, with very scarce biotite and feldspars. A secondary schistosity is well-defined by alternation of quartz-rich and muscovite layers. The small muscovite grains show characteristic shapes and absence of intra-crystalline deformation. A representative sample of these quartzites (sample 101, Fig. 2), collected immediately beneath the main thrust, has a muscovite age of 395.7 + 0.2 Ma, similar age to that of a mylonitic gneiss of the hanging wall, in the central part of the range (sample 105). The granitic dykes at Cerro Valdivia, emplaced in meta-pyroxenites, show conspicuous schistosity. Sample 110 (Fig. 2), contains anhedral quartz crystals along bands of deformation associated with sub-idiomorphic crystals of biotite and muscovite. The muscovite age obtained in this granitic dyke yielded 425.0 + 0.2 Ma, within the age range of the previous samples. The hornblende spectra obtained in the basic and ultrabasic metamorphic rocks of the Pie de Palo complex have two groups. The first group, represented by geologically significant values in amphibolites, yielded plateau ages of 463.1 + 0.9 (104), 360.4 4- 0.8 (108) and 423.7 + 0.9 Ma (109). The second group has older apparent ages, without plateau. However, it could be geologically significant that these ages increase in the high temperature increments, approaching 1050Ma in the meta-pyroxenite of Cerro Valdivia (sample 111). These ages are similar to the 1060-1080 Ma obtained by U-Pb in zircons from amphibolites of this belt by McDonough et al. (1993). The 4~ hornblende age obtained from an amphibolite immediately above the Pirquitas thrust (sample 104) is older than the muscovite ages. This is consistent with their different isotopic closure temperatures.

155

Discussion The 4~ ages presented here are related to the cooling and uplift of a complex deformational episode superimposed on the Mid-Proterozoic magmatic crystallization and metamorphism of the Pie de Palo complex. In order to evaluate the tectonic significance of these episodes the tectonic events seen in the adjacent eastern Precordillera must be considered. Since the early studies of Furque (1972) and Baldis et al. (1982), a series of diastrophic events has been recognized in the eastern Precordillera. Mid-Late Ordovician to Early Silurian discontinuities represent tectonic events whose timing is constrained by an abundant fossil record (Astini et al. 1995, 1996), and may be converted to numerical ages using the time-scale of Gradstein & Ogg (1996): (1) beginning of collision and drowning of the Precordillera carbonate platform against Gondwana (Llanvirn-early Llandeilo, 470-460Ma); (2) normal faulting by flexural extension of the Precordillera associated with tectonic loading of the Gondwana margin and (3) development of the first foreland basin (Late Caradoc-Llandovery, 450-430 Ma). These episodes were directly related to the collision between the Cuyania composite terrane and the proto-margin of Gondwana. In the Pie de Palo basement these deformational events are represented by ductile shear reactivation of the Precambrian fabrics associated with thrusting and subsequent uplift, whose timing may be related to some of the 4~ ages obtained in this study. Subsequent uplift of the eastern Precordillera occurred during the Early to Middle Devonian. This resulted in cannibalization of the former foreland basin and the development of the Punta Negra foreland basin further west during the Early to Middle Devonian (410380Ma) (Astini 1996). 4~ ages from muscovite concentrates within this time interval thus appear to record the last pre-Andean uplift of the Pie de Palo region. This episode is linked to the earliest stage of the docking of Chilenia to the western margin of Precordillera (Ramos et al. 1984). Correlation of these events with the cooling ages of the Sierra de Pie de Palo shows that thrusting and subsequent uplift correlate with the ductile deformational episodes. These are represented by mylonites and protomylonites, affecting the crystalline basement with a dominant west-vergent transport. A widely distributed episode close to 395Ma, detected in the new 4~ ages, as well as in previous K-Ar

156

V. A. RAMOS E T AL.

dating (Linares & Gonzfilez 1990), indicates the cessation of ductile deformation in the Pie de Palo region. The suture between the Cuyania terrane and the proto-margin of Gondwana should be located to the east of the Sierra de Pie de Palo. A large mylonitic belt observed along the western border of Sierra de Valle F6rtil, east of the Bermejo river (Fig. 1), is one of the best candidates for this suture. Deep seismic reflection data reveals a striking discontinuity in the crustal structure of either sides of this belt between the Cuyania and Pampia terranes (Zapata 1996). These deformational episodes may be compared with equivalent deformation that led to similar time-breaks in the Appalachians, recently recorded by 4~ plateau ages (Castonguay et al. 1997). The sequence of continuous events between Mid-Ordovician (Taconian deformation) and Mid-Devonian (Acadian deformation) times are comparable to the age record of the Sierra de Pie de Palo. But, the parallel evolutionary paths between these two orogens are better explained by coeval processes controlled by early Palaeozoic global mobility, resulting in simultaneous tectonic events in far-separated geographical areas (Astini et al. 1995, 1996), rather than by a continuous unified belt as previously proposed (Dalla Salda et al. 1992a, b; Dalziel et al. 1994).

Concluding remarks The collision between the Cuyania composite terrane and the proto-margin of Gondwana began in Mid-Ordovician time (c. 470 Ma) and its completion is constrained by 4~ muscovite ages ranging down to 394Ma. The distribution in time and space of these ages relates to a complex belt of ductile deformation, thrusts and overthrusts affecting the Pie de Palo basement. The westward vergence of the thrusts, as well as the fossil record in the fill of adjacent foreland basins, show a coincidence between the timing of basin evolution and periods of thrusting and uplift derived from 4~ plateau ages. This correlation is consistent with the hypothesis that the Precordillera, as part of the Cuyania composite terrane, docked against the protomargin of Gondwana in Mid-Ordovician time, but the effects of deformation and uplift, with subsequent cooling lasted until the Early Devonian. The last episodes of deformation were related to the beginning of the docking of Chilenia further to the west.

The present research was funded by UBACYT Ex-132 grant and CONICET PIP 4162/96. This is a contribution to IGCP 376 Laurentia Gondwana Connections before Pangea.

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Hidrocarburos,

Buenos Aires, Actas, I, 541-550. CASTONGUAY, S., TREMBLAY, A., RUFFET, G,, FERAUD, G., PINEY, N. & SOSSON, M. 1997. Ordovician and Silurian metamorphic cooling ages along the Laurentian margin of the Quebec Appalachians: bridging the gap between New England and Newfoundland. Geology, 25, 583-586.

EARLY PALAEOZOIC PRECORDILLERA DOCKING CLIFF, R. A. 1985. Isotope dating in metamorphic belts. Journal of the Geological Society, London, 142, 97-110. CRIADO ROQUE, P. 1972. Bloque de San Rafael. In: LEANZA, A. (ed.) Geologia Regional Argentina. Academia Nacional de Ciencias, C6rdoba, 283-295. DALLA SALDA, L. & VARELA, R. 1984. E1 metamorfismo en el tercio sur de la Sierra de Pie de Palo, San Juan. Revista de la Asociacidn Geoldgica Argentina, 39, 68-93. --, CINGOLANI, C. & VARELA, R. 1992a. Early Paleozoic orogenic belt of the Andes in southwestern South America: Result of LaurentiaGondwana collision? Geology, 20, 617-620. - - , DALZIEL,I. W. D., CINGOLANI,C. A. & VARELA,R. 1992b. Did the Taconic Appalachians continue into southern South America? Geology, 20, 1059-1062. DALLMEYER, R. D. & GIL-IBARGUCHI, I. 1990. Age of amphibolitic metamorphism in the opholitic unit of the Morais allochthon (Portugal): Implications for early Hercynian orogenesis in the Iberian Massif. Journal of the Geological Society, London, 148, 873-878. DALZIEL, I. W. D., DALLA SALDA, L., CINGOLANI, C. & PALMER, P. 1996. The Argentine Precordillera: A Laurentian terrane? GSA Today, 6, 16-18. , & GAHAGAN, L. M. 1994. Paleozoic Laurentia-Gondwana interaction and the origin of the Appalachian-Andean mountain system. Geological Society of America Bulletin, 106, 243-252. FURQUE, G. 1972. Los movimientos caled6nicos en Argentina. Museo de La Plata, Revista (Nueva Serie) VIII, Geologla, 65, 129-136. GRADSTEIN, F. & OGG, J. 1996. A Phanerozoic time scale. Episodes, 19, 3-5. HARRISON, T. M. 1981. Diffussion of 4~ in hornblende and its consequences for 4~ age spectra and closure temperature. Geochimica et Cosmochimica Acta, 50, 247-253. HEREDIA, S. 1996. E10rdovicico del arroyo Pondn Trehue, sur de la provincia de Mendoza. XIII Congreso Geol6gico Argentino y III Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, I, 601-605. HOLMBERG, E. 1973. Descripci6n geoldgica de la Hoja 29d Cerro Nevado, Provincia de Mendoza. Servicio Nacional Minero Geolfgico, Boletin, 144, 1-71. JORDAN, T., ZEITLER, P., RAMOS, V. A. & GLEADOW, A. J. W. 1989. Thermochronometric data on the development of the basement peneplain in the Sierras Pampeanas, Argentina. Journal of South American Earth Sciences, 2, 207-222. KAY, S. M., ORRELL, S. & ABBRUZZI, J. M. 1996. Zircon and whole rock Nd-Pb isotopic evidence for a Grenville age and a Laurentian origin for the Precordillera terrane in Argentina. Journal of Geology, 104, 637 648. LEVERATTO, M. A. 1968. Geologia de la zona al oeste de Ulhin-Zonda, borde oriental de la Precordillera de San Juan, eruptividad subvolcfinica y estructura. Revista de la Asociaci6n Geol6gica Argentina, 23, 129 157.

157

1976. Edad de intrusivos cenozoicos en la Precordillera de San Juan y su implicancia estratigrfifica. Revista de la Asociacidn Geol6gica Argentina, 31, 53-58. LINARES, E. & GONZAkEZ, R. R. 1990. Catfilogo de edades radiom&ricas de la Repflblica Argentina 1957-1987. Asociacidn Geol@ica Argentina, Publicaciones Especiales Serie B, Did6ctica y Complementaria, 19, 1-628. - - , LLAMBiAS,E. & LATORRE, C. O. 1980. Geologia de la provincia de La Pampa, Repfiblica Argentina y geocronologia de sus rocas metam6rficas y eruptivas. Revista de la Asociaci6n Geol6gica Argentina, 35, 87-146. McDoNou6H, M. R., RAMOS, V. A., ISACHSEN, C. E., BOWRING, S. A. & VUJOVICH, G. I. 1993. Nuevas edades de circones del basamento de la sierra de Pie de Palo, Sierras Pampeanas Occidentales de San Juan: sus implicancias para los modelos del supercontinente proterozoico de Rodinia. XIII Congreso Geol6gico Argentino y 111 Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, III, 340-342. MCEACHERN, S. J., HANMER, S. & VAN BREEMEN, O. 1990. Geochronology and structure of the Central Metasedimentary belt boundary thrust zone, Pembroke, Ontario. Geological Association Canada, Abstracts with Program, A87. Nt)~qEZ, E. 1979. Descripci6n geol6gica de la hoja 28d Estaci6n Soitu6, provincia de Mendoza. Servicio Geol6gico Nacional, Boletln, 166, 1-67. Pucc1, I. C. 1976. Contribuci6n al conocimiento del subsuelo de General Alvear (provincia de Mendoza) y regiones adyacentes. PhD thesis 1529, Universidad de Buenos Aires, Buenos Aires. RAMOS, V. A. 1995. Sudam6rica: un mosaico de continentes y oc6anos. Ciencia Hoy, 6, 24-29. -& VUJOVICH, G. 1995. Hoja Geol6gica San Juan, escala 1: 250.000. Direcci6n Nacional del Servicio Geol6gico, Subsecretaria de Mineria (open-file report). - - , JORDAN, T., ALLMENDINGER,R. W., KAY, S. M., CORTES, J. M. & PALMA, M. A. 1984. Chilenia un terreno aldctono en la evolucidn paleozoica de los Andes Centrales. IX Congreso Geol6gico Argentino, Buenos Aires, Actas, II, 84-106. , MPODOZIS, C., KAY, S. M., CORTES, J. M. ~: PALMA, M. A. 1986. Paleozoic terranes of the Central Argentine-Chilean Andes. Tectonics, 5, 855-880. SAMPSON, S. D. & ALEXANDER,E. C. 1987. Calculation of the interlaboratory 4~ dating standard, MMhb-1. Chemical Geology (Geoscience Section), 66, 500-506. TOSDAL, R. M. 1996. The Amazon-Laurentian connection as viewed from the Middle Proterozoic rocks in the central Andes, western Bolivia and northern Chile. Tectonics, 15, 827-842. , MUNIZAGA, F., WILLIAMS, W. C. & BETTENCOURT, J. S. 1994. Middle Proterozoic crystalline basement in the Central Andes, Western Bolivia and Northern Chile: a U-Pb and Pb isotopic perspective. 7~ Congreso Geol6gico Chileno, Concepcidn, Actas, II, 1464-1467.

158

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V. A. R A M O S E T AL.

VuJovIcH, G. & KAY, S. M. 1996. Evidencias geoquimicas del origen y ambiente geol6gico de las rocas metam6rficas de composici6n mfifica e intermedia de las Sierras Pampeanas Occidentales. XIII Congreso Geol6gico Argentino y III Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, V, 273~91. & --1998. A Laurentian? Grenville age oceanic arc/back-arc terrane in the Sierra de Pie Palo, Western Sierras Pampeanas, Argentina. This volume. & RAMOS, V. A. 1994. La Faja de Angaco y su relaci6n con las Sierras Pampeanas Occidentales, -

Argentina. 7~ Congreso Geol6gico Chileno, Concepci6n, Actas, I, 215-219. WICHMANN, R. 1928. Contribuci6n a la geologia de los departamentos Chicalc6 y Puel6n de la parte occidental de la Pampa Central. In: Direcci6n General de Minerla, Geologia e Hidrogeologla, Publicaci6n, 40, 3-14. YORK, D. 1969. Least square fitting of a straight line with correlated errors. Earth Science and Planetar)' Letters, 5, 320-324. ZAPATA, T. R. 1996. Crustal evolution of the Precordillera thrust belt, Bermejo basin, Argentina. PhD thesis, Cornell University, Ithaca.

A Laurentian? Grenville-age oceanic arc/back-arc terrane in the Sierra de Pie de Palo, Western Sierras Pampeanas, Argentina GRACIELA

I. V U J O V I C H 1 & S U Z A N N E

MAHLBURG

KAY 2

I Laboratorio de Tect6nica Andina, Departamento de Ciencias Geol@icas, Universidad de Buenos Aires-CONICET, Argentina. Pabell6n 2, Ciudad Universitaria (1428), Buenos A ires, Argentina (e-mail: [email protected]) 2Department of Geological Sciences, Snee Hall, Cornell University, Ithaca, NY, 14853, USA (e-mail. [email protected]) Abstract" Whole-rock trace element, field, and petrographical studies indicate that the

protoliths of highly deformed metamorphic rocks of the Pie de Palo Complex, Argentina, formed in an oceanic setting like that of SW Pacific supra-subduction complexes (e.g., Lau Basin). Age constraints suggest that this Mesoproterozoic complex, which includes some of the most primitive Grenville-age magmas known on Earth, formed in <60 million years. Amphibolites (Las Pirquitas/Quemado region) with N-MORB like light-depleted REE patterns and Hf/Ta/Th relations, associated with Fe-rich schists interpreted as hydrothermally altered lavas and gabbros, formed in a back-arc spreading centre. Other ultramafic/mafic units (Quebrada del Gato, Cerros Valdivia/Barboza) formed as arc cumulates, and (central region) as ridge or arc/back-arc lavas (including shoshonite). Intermediate and silicic schists/gneisses are interpreted as arc region greywackes and silicic lavas, and marbles as reef deposits. Quebrada del Gato gneisses with trace elements indicating high pressure residual mineral assemblages are interpreted as younger subduction-related magmas emplaced in a thickened crust. The Pie de Palo Complex probably formed offshore of eastern Laurentia, and subsequently collided with Laurentia east of the Llano uplift, Texas, in Grenvillian times. This block, along with the Precordillera, was rifted flom Laurentia in the latest Precambrian, arriving in Gondwana by Ordoviclan time.

The nature of the Mesoproterozoic to Palaeozoic basement of the Sierras Pampeanas in central west Argentina is important in unravelling the Mesoproterozoic to Early Palaeozoic history of terrane interactions between Laurentia and Gondwana. Particularly important is the basement of the Sierra de Pie de Palo that lies just east of the Precordillera Range. The Precordillera is widely accepted as a terrane, or part of a terrane, that separated from Laurentia in the late Precambrian to Cambrian and was accreted to Gondwana in the Ordovician (Fig. 1; e.g., Dalziel et al. 1996). The Sierras Pampeanas themselves are a series of crystalline basement blocks, whose most recent uplift has been along high angle reverse faults produced in association with Neogene shallowing of the Andean subduction zone (e.g., Jordan & Allmendinger 1986). Much of the outcrop in the Sierra de Pie de Palo, as well as in the Cerros Barboza and Valdivia blocks to the southwest, is composed of the Mesoproterozoic Pie de Palo Complex (Fig. 2) as defined by Ramos & Vujovich (in press). This complex, which is principally made up of greenschist to amphibolite facies schists and gneisses, includes a belt of Alpine-type mafic-ultramafic rocks (Villar 1985). Discussion

of various aspects of these units can be found in Llano & Escalante (1979), Dalla Salda & Varela (1982, 1984), Castro de Machuca (1981), Ramos et al. (1993), Castro de Machuca & Conte Grand (1994), Castro de Machuca et al. (1995), Vujovich & Ramos (1994), Vujovich & Kay (1996), and Ramos & Vujovich (in press). The purpose of this paper is to present new data and a synthesis of the chemical characteristics of a reconnaissance field study of the various units. The data show that the Pie de Palo Complex contains one of the best-exposed, near-juvenile Mesoproterozoic oceanic arc/back-arc complexes on Earth. Possible modern analogues for the protolith of the Pie de Palo Complex are supra-subduction complexes in the southwestern Pacific Ocean. The Lau Basin region is used for comparison. The field, petrological and geochemical characteristics of the Pie de Palo Complex provide clues for finding correlative units in Laurentia. As the Precambrian rocks from the western Sierras Pampeanas were metamorphosed under greenschist to amphibolite facies conditions and are extremely deformed, their primary magmatic characteristics are best deciphered through the chemistry of their immobile elements, particularly Ti, the REEs, Ta, Th, Sc, Hf, Cr and Ni.

VUJOVICH,G. I. & KAY, S. M. 1998. A Laurentian? Grenville-age oceanic arc/back-arc terrane in the Sierra de Pie de Palo, Western Sierras Pampeanas, Argentina. In: PANKHURST,R. J. & RAPELA,C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 159-179.

160

G. I. VUJOVICH & S. M. KAY

Distribution, lithology, and age of the Pie de Pale Complex

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' deposits Palaeozoic sedimentary deposits Cambro-Ordovician platform sequences Early Palaeozoic granitoids Proterozoic-Palaeozoic metamorphic rocks

Fig. 1. Map of the western Sierras Pampeanas and Precordillera region over the modern shallow subduction zone in the central Andes, showing the location of the Sierra de Pie de Palo, and a possible suture zone between the Cuyania and Pampia terranes (e.g., Ramos et al. this volume). The Cuyania terrane incorporates the Precordillera terrane with the Pie de Palo region.

Other elements, such as the alkalis (Na, K, Cs), alkaline earths (Ba, Sr) and in some cases U, have to be dealt with carefully as they are susceptible to alteration during weathering and metamorphism. Recognizing these limitations, the procedure here is to use diagnostic major and trace elements qualitatively to interpret the origin and tectonic environment of the protoliths of these deformed metamorphic rocks. Differentiating between igneous and juvenile sedimentary protoliths is at times difficult, if not impossible. Some mobile elements, such as Na and Sr, are useful for this purpose. Deciphering arc/back-arc environments from intraplate and MORB sources is based on overall characteristics, but is particularly dependent on N-MORB and chondrite-normalized trace element patterns. High normalized concentrations of La relative to Ta (La/Ta ratios >25) are taken to be important signatures of arc/back-arc protoliths. The use of chemical characteristics to differentiate magmatic and tectonic environments is discussed in general references (e.g., Wilson 1989; Rollinson 1993; Taylor & McLennan 1985), and is not covered in detail here.

A general map of the Sierra de Pie de Palo illustrating the outcrop region of the Pie de Palo Complex is shown in Fig. 2. The grade of metamorphism is principally in the greenschist and amphibolite facies, although Dalla Salda & Varela (1984) report granulite-facies metamorphism in the southernmost part of the range. In general, the central part of the Pie de Palo Complex is composed of biotite-muscovitegarnet-oligoclase-bearing schists and gneisses, the eastern part by biotite-garnet-feldspar-bearing schists and gneisses, and the western part by a mafic-ultramafic complex. Associated with the schist and gneisses of the central and eastern regions are bodies of massive amphibolite and amphibolitic schists. Also present among the gneisses are zones of incipient migmatite, characterized by veins and lenses rich in K-feldspar and quartz. The protoliths of the schists and gneisses have been largely interpreted as silicic magmas and immature greywackes (Ramos et al. 1993; Vujovich & Kay 1996). The protoliths of the mafic-ultramafic belt have been interpreted as ultramafic and mafic cumulates and flows (e.g., Ramos et al. 1993; Castro de Machuca & Conte Grand 1994; Vujovich & Kay 1996). Samples in this study come from the Cerros Valdivia and Barboza, the Quebradas del Gato, El Quemado, Las Pirquitas, Guayaupa-Lima, the Mogote Corallitos region, and the Nikizanga block (see Fig. 2). Ultramafic and mafic rocks are particularly important in understanding the origin of the Pie de Palo Complex. Rock types include serpentinite, metagabbro, metadiorite, amphibolite, and a variety of mafic schists (e.g., Dalla Salda & Varela 1984; Ramos & Vujovich in press). The majority of the outcrops are chlorite, talcchlorite and amphibolite schists with nematoblastic to decussate textures. The most important minerals are tremolite-actinolite, talc, chlorite, albite, and various varieties of epidote (zoisite). Effects of cataclasis and proto-mylonitization are pervasive. Most of these units occur in a belt that strikes NNE and extends for c. 20 kilometres between the Quebradas (canyons) Las Pirquitas and Piedras Pintadas on the western side of the Sierra de Pie de Palo (Fig. 2). Sequences of these units are repeated in a series of NNE-striking, ENE-dipping thrust faults above the principal Las Pirquitas fault. This main fault carries the Pie de Palo Complex over the siliciclastic and carbonate platform units of the Caucete Group, whose age is unknown (see Ramos et al. 1996). Mafic and ultramafic

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R b - S r isochron in Varela & Dalla Salda (1993). Zircon ages from central region biotitemuscovite-garnet and biotite-garnet gneisses between the Mogote Corralitos and the Quebrada Grande (see Fig. 2) range from 1090 to

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GRENVILLE OCEANIC TERRANE IN PIE DE PALO 1060 Ma. Gneisses similar to these extend northward into the Quebradas de Guayaupa-Lima region investigated in this study. Zircons from meta-rhyolite cropping out in the headwaters of the Quebrada del Molle gave ages of c. 1080 Ma. Metamorphic zircons from amphibolites in the central part of the Quebrada del Molle near the mafic-ultramafic belt gave ages of 1066 Ma. A Rb-Sr reference isochron age of 1028Ma from gneisses and granites in the southern part of the Pie de Palo Complex also indicates a Grenville age. A Grenville age for the mafic-ultramafic belt is based on a zircon fraction separated from a biotite-amphibole-garnet schist in Cerro Barboza that gives a 2~176 age of 1118 Ma (Alberto Saal pers. comm. 1997). This schist, whose protolith is considered to have been a greywacke, is intercalated with amphibolite in the mafic-ultramafic belt. Other evidence for a Grenville age comes from complex 4~ release spectra on hornblende separates from mafic-ultramafic belt samples in the southern Sierra de Pie de Palo near the Caucete cemetery (Fig. 2) and Cerro Valdivia in Ramos et al. (this volume). These authors argue that high temperature increments from the Cerro Valdivia hornblende separate indicate increasingly older ages that approach 1050 Ma. Grenville-age basement has also been reported in the Precordillera to the west (Abbruzzi et al. 1993; Kay et al. 1996). This basement is not exposed, but is known from xenoliths in the Miocene Cerro Blanco-Ullfin volcanic complex (Leveratto 1968). Xenolith types include feldspathic gneiss, and mafic pyroxene and pyroxenegarnet granulite in Cerro de Los Bafios, and amphibolite and greenschist facies mafic schist and acidic gneiss at a second locality along the valley just to the north (Abbruzzi et al. 1993). Evidence for a Grenville age comes from U-Pb zircon upper intersect ages of 1102 + 6 Ma for a mafic schist and a less well constrained age of 1096+ 50Ma for an acidic xenolith (Kay et al. 1996). Abbruzzi et al. (1993) and Kay et al. (1996) use chemical data to argue for an oceanic arc,/ back-arc environment near a continental margin for the protolith of these xenoliths.

Chemical characteristics and protoliths of the Pie de Palo Complex Lithological units of the Sierra de Pie de Palo region are considered in four groups based on their geographical distribution and geochemical characteristics. They are: (1) western belt ultramafic/mafic units with arc characteristics from

163

the Quebrada del Gato and Cerros Valdivia and Barboza, (2) western belt mafic units with spreading centre characteristics from the Quebradas Las Pirquitas and E1 Quemado, (3) mafic units from the Quebradas Guayaupa-Lima and Molle, the Nikizanga block, and the Mogote Corralitos region within the central and eastern schist-gneiss complex, and (4) intermediate and silicic gneisses from the Quebrada GuayaupaLima, the Nikizanga block, and the Quebrada del Gato. Chemical analyses are given in Tables 1 to 4. Trace element data are plotted in Figs 3 to 7.

A n arc environment. mafic and ultramafic cumulates f r o m the Quebradas del Gato and Cerros Valdivia and Barboza

Following Vujovich & Kay (1996), petrological and geochemical characteristics are used to classify the protoliths of the mafic-ultramafic group in the Quebrada del Gato and the Cerros Valdivia and Barboza into two arc-derived petrogenetic facies: ultramafic and mafic cumulates. Detailed petrographical descriptions of these units can be found in Ramos & Vujovich (1998). The ultramafic units. The ultramafic group is composed primarily of serpentinite and metapyroxenite bodies. The most important serpentinite outcrops occur in the Quebradas del Gato and Piedras Pintadas where massive to schistose lenticular bodies are interleaved with greenschist-facies mafic schists. The mineralogy of these serpentinites is now principally antigorite, talc and tremolite. Abundant opaque minerals, mainly magnetite and chromite, also occur (Castro de Machuca et al. 1995). Metapyroxenite bodies in Cerro Valdivia are principally composed of relict uralitized pyroxene, amphibole, chlorite and opaque minerals, with minor apatite and titanite. Chemical analyses of two typical tremolitebearing serpentine schists, Quebrada del Gato samples PP96 and 97a, are shown in Table 1 and Fig. 3a b. Support for an ultramafic cumulate protolith comes from the presence of chromite, and petrographical evidence for relict olivine and pyroxene grains. The combination of high Ni (c. 1800ppm) and Cr (c. 4700ppm), and low Sc and REE concentrations is consistent with the original mineralogy being dominated by olivine and chromite. The combination of Eu anomalies, high Sr (460-472ppm) and low Na concentrations might indicate the original presence of Ca-rich plagioclase. However, secondary mobility of Na, Eu, and Sr in fluids is likely in these

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Nikizanga Block PP25

SiO2 TiO2 A1203 FeO MnO MgO CaO

65.78 0.60 17.29 4.80 0.15 1.73 2.31 4.19 2.83 0.03 99.71

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Blank in trace element column indicates concentration below level of detection La 45.1 48.3 31.0 35.6 46.9 11.1 Ce 106.1 105.8 69.0 78.3 103.9 21.7 Nd 55.0 41.8 33.6 37.4 45.8 9.07 Sm 11.2 8.47 7.58 7.91 10.1 2.19 Eu 2.03 0.970 1.48 1.40 1.73 0.828 Tb 1.61 0.930 1.10 1.04 1.32 0.301 Yb 5.65 3.37 4.11 3.33 4.96 0.928 Lu 0.786 0.465 0.593 0.468 0.711 0.120 Sr 257 206 192 180 191 542 Ba 749 692 498 670 1445 267 Cs 3.3 1.7 2.2 1.5 191 1.0 U 4.0 2.8 1.8 1.7 2.5 0.6 Th 13.9 12.5 7.5 8.9 12.1 1.7 Hf 14.3 8.3 7.5 7.6 10.3 2.3 Ta 1.4 0.9 1.1 0.7 1.1 0.11 Sc 10.0 6.9 16.8 9.4 11.1 12.3 Cr 40 9 83 36 42 11 Ni 26 10 53 41 26 9 Co 11 5 21 10 11 12

13.2 26.2 10.5 2.09 0.604 0.200 0.616 0.087 446 265 1.0 0.9 2.7 2.1 0.12 8.5 20 13 9

19.5 36.6 13.2 2.31 0.593 0.179 0.439 0.042 416 518 0.8 2.8 6.7 3.1 0.24 4.1 29 22 6

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highly altered rocks. The presence of chromite requires that the p r i m a r y m a g m a s that generated these cumulates were derived from a very depleted mantle source as expected in an oceanic arc or back-arc setting. Very low Ta concentrations ( < 0 . 0 5 p p m ) indicate that the primary m a g m a had a high, characteristically arc-like L a / T a ratio. Analyses of Cerro Valdivia samples V42 and V43, shown in Table 1 and Fig. 3c-d, are consistent with ultramafic cumulate protoliths in which the principal minerals were orthopyroxene and clinopyroxene (47-54 p p m Sc), olivine

18.8 4.5 10.7 0.59 4 49 5.3

30.8 4.7 9.4 0.57 4 42 4.8

24.0 5.1 12.0 1.24 7 106 2.8

(403 p p m Ni), and Cr spinel (2462-3172 p p m Cr). The original presence of both pyroxenes is consistent with CaO concentrations of 13-14% (a clinopyroxenite has c. 20% CaO). A scarcity of primary plagioclase is indicated by Sr concentrations of 257-309 ppm, and by m u c h of the 3.44.5% AlzO3 being concentrated in the spinel. Plagioclase should have been m o r e a b u n d a n t in V43 than V42 (higher Sr and A1203 concentrations, positive Eu anomaly). The overall shapes of the R E E patterns, particularly that of V43, are unlike that of clinopyroxene, the only phase with a significant R E E concentration, and must

GRENVILLE OCEANIC TERRANE IN PIE DE PALO 100 50

Quebrada del Gato Cumulates PPIO1 ~ / PP97b

,~ 10 ~

a)

167

100 Quebrada del Gato

b)

Cumulates

/~

PP102

-

" l!..... 0.5

PPlO0

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~~ 0.1

g/

PP97a

0.1

~

\

/

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~ ~

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Cerros Barboza & Valdivia A _~ \

o B23 x v45 ~ V43 9 V44

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--

0.5 0.1 La Ce

Nd

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Tb

o B23 V45 ~ 9 V43 9 V44 " V42

YbLu

0.1 Cs

Ba ~/" U Th

La Ta

Sr Ce

Hf Nd

Eu Sm

Yb Tb

Lu

3 . Chondrite-normalized rare earth element (REE) and N-MORB normalized extended trace element patterns for Pie de Palo Complex western ultramafic belt/mafic samples, from the Quebrada del Gato in the Sierra de Pie de Palo (a and b) and the Cerros Barboza and Valdivia (c and d). Protoliths are interpreted as ultramafic and mafic cumulates in an arc/back-arc setting. See text for discussion. Data from Table 1. Normalization in the REE plots is to the Leedy Chondrite [La (0.378), Ce (0.976), Nd (0.716), Sm (0.23), Eu (0.0866), Tb (0.0589), Yb (0.249) and Lu (0. 0383)] and in the extended element plot to the N-MORB composition in Sun & McDonough (1989).

Fig.

mirror the intercumulus liquid. The pattern of V43, like those of the Valdivia amphibolites discussed below, supports an arc origin.

The mafic cumulates. Massive metagabbro, metadiorite, amphibolite, and garnet amphibolite bodies are present along with schistose units in the Quebrada del Gato and in Cerros Barboza and Valdivia. These bodies show compositional and grain-size layering. Some have granular textures, whereas others have cumulate textures. All exhibit some degree of cataclasis. The metagabbro bodies are principally composed of saussuritized plagioclase and relict pyroxene altered to amphibole, and the metadiorite bodies of aggregates of saussuritized plagioclase and clinozoisite. Amphibolite bodies are dark green, massive units principally composed of amphibole, saussuritized plagioclase, epidote and minor quartz. Some have biotite and/or garnet. Amphibolite samples interpreted as gabbroic cumulates from the Quebrada del Gato include PP100, PP102, PP97b, and PP101 in Fig. 3a-b and Table 1. An arc protolith for all of these samples is consistent with evidence for primary

amphibole and F e - T i oxides, and the behaviour of Ta relative to the light REE. In detail, the protolith of PP100 is best explained as a gabbro with an original mineralogy of plagioclase, evidenced by a large positive Eu anomaly (Eu/Eu* = 2.59), pyroxene as evidenced by light REE depletion and high Sc (46ppm) and Cr (98 ppm) concentrations, and minor olivine as evidenced by the Ni (45ppm) concentration. A very high La/Ta ratio (102) is consistent with arc derivation. The characteristics of garnet amphibolite PP102, which include an Eu/Eu* ratio of 2.96, indicate a plagioclase-rich gabbroic cumulate. A high FeO concentration indicates Fe-rich phases, probably amphibole and titanomagnetite. The concave upward shape of the REE pattern is best explained by removal of amphibole. The overall trace element composition is dominated by the pattern of an intercumulus magma with high H f (30ppm), U, and Th concentrations, and an arc-like La/Ta ratio (31). Samples PP101 and PP97b have convex upward REE patterns indicative of amphibole accumulation. A positive Eu anomaly ( E u / E u * = 1.4) in PP97b signals cumulus plagioclase. High FeO concentrations (12-15 %)

168

G. I. VUJOVICH & S. M. KAY

100

50

....

a) 100

Quebradas Las ~/r,r~u.Las, D~ ,-, El Quemado Amphibolites

Quebradas Las Pirquitas/ El Quemado Amphibolites

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0.1 100

100 50

Quebrada Las Pirquitas Gabbros

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Quebrada Las Pirquitas Mafic Schists

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.

.

.

.

.

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.

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Quebrada Las Pirquitas Mafic Schists

1

0.5

1

0.1 La Ce

Nd

Sm Eu

Tb

YbLu

0.I L Cs

v

Ba

U Th

La Ta

Sr Ce

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Eu Sm

Yb Tb

Lu

Fig. 4. Chondrite-normalized REE and N-MORB-normalized extended trace element patterns for Pie de Palo Complex western ultramafic/mafic belt samples from the Quebradas Las Pirquitas and El Quemado. Protoliths are interpreted as forming in a rift setting in a back-arc basin. (a) and (b) Mafic amphibolites interpreted as spreading centre basaltic dykes and lavas. (c) and (d) Metagabbros interpreted as cumulate gabbros. (e) and (f) Fe-rich mafic schists interpreted as hydrothermally altered rift-related magmatic rocks, possibly with ferromanganese crusts. Further discussion in text. Data in Table 2. Normalization as in Fig. 3. in both samples reflect amphibole and titanomagnetite. A high, non-arc like Ta concentration (Fig. 3b) and La/Ta ratio (6) in PP97b can be explained by cumulate amphibole and titanomagnetite, both of which incorporate Ta. Barboza amphibolite sample B23 (Fig. 3c-d) can be interpreted as an original arc-derived amphibole-plagioclase-Fe-Ti oxide-bearing cumulate. Support comes from the whole-rock composition which mimics a Ti-rich amphibole with Fe-Ti oxide (4.9% TiO2; 17.9 FeO), and a high Sr content (858ppm) and a large positive Eu anomaly (Eu/Eu* = 1.6) which indicate

plagioclase. Further evidence comes from an amphibole-like concave downward R E E pattern (Fig. 3c). As with PP97b, a somewhat low La/Ta ratio (22) for an arc magma can be explained by accumulation of amphibole and FeTi oxide phases. Amphibolite samples V44 and V45 from Cerro Valdivia (Fig. 3c-d) are also best interpreted as having arc cumulate protoliths. Both have whole-rock major element compositions mimicking amphibole, consistent with an amphibole-bearing gabbroic cumulate, or less likely, a phenocryst-rich lava. The importance of

GRENVILLE OCEANIC TERRANE IN PIE DE PALO 500 500~__~pp10 CentralPie de Palo a) 10( 100 Shoshonite - - - ~ ~ ~ p h i b o l i t e s 5C 5 0 ~ _ ~ _ 10 5 1

~

1-1

169

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PP13

.

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~k

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Quebrada Guayaupa Mafic Units

0.5 0.1

PP92

Nd

Sm Eu

Tb

0.1 - Ba

YbLu

Cs

U

Th

La

Ya

Sr

Ce

Hf

Nd

Eu

Sm

Yb-

Tb

Lu

Fig. 5. Chondrite-normalized REE and N-MORB-normalized extended trace element patterns for Pie de Palo Complex mafic amphibolite bodies intercalated with schists and gneisses in the central and eastern parts of the Sierra de Pie de Palo. Protoliths are interpreted as being arc/back-arc and rift-related. (a) and (b) Amphibolites interpreted as arc/back-arc lavas, PP10 has shoshonitic characteristics. (e) and (d) Amphibolites from the Quebrada del Guayaupa interpreted as back-arc rift-related (PP62 and PP73), and arc-related (PP85, PP92, and PP79) lavas. PP79 could be an immature arc-derived sediment. U and Sr spikes in P73 attributed to interaction with carbonate units. Further discussion in text. Data in Table 3. Normalization as in Fig. 3.

plagioclase is indicated by positive Eu anomalies ( E u / E u * = 1.17-1.29) and high Sr concentrations (847-859ppm), of clinopyroxene by high Cr (584-916ppm) and Sc (32 34ppm) concentrations, and of olivine by high Ni (130142ppm) concentrations. The similarity of the chondrite-normalized REE and N-MORBnormalized trace element patterns of V44 and V45 to that of clinopyroxenite V43 argues that these patterns are all largely controlled by intercumulus liquid. Extreme La/Ta ratios (>100) support an arc environment, as does the association of these units with ultramafic cumulates whose protoliths are interpreted as chromite-bearing clinopyroxenite (V42 and V43). Low Th and high Ba, Cs, and U concentrations (Fig. 3b-d) suggest a depleted and subsequently hydrothermally-enriched arc source.

A spreading ridge environment. the mafic sequence at the Quebradas el Quemado and Las Pirquitas A different sequence of metamorphic rocks characterizes the mafic-ultramafic belt in the Quebradas el Quemado and Las Pirquitas

regions. These units, which crop out just above the Las Pirquitas thrust, are generally upper greenschist facies mafic amphibolites associated with metagabbros, metadiorites, micaceous schists, and opaque-rich, amphibole-talc schists. The entire sequence has been subjected to intense ductile deformation. The rocks are proto-mylonites and mylonites occurring in tight isoclinal, overturned and refolded folds, and imbricated thrusts above the east-verging principal thrust. The thrust contact itself is characterized by interdigitation of Caucete Group quartzite and quartz-muscovite schists with the mafic rocks, on both a microscopic and macroscopic scale. Samples collected near the principal fault consist of interdigitated bands of quartz-mica, mica, and talc-amphibole schists. Some have layers of biotite and muscovite interspersed with lenses of quartz, rounded plagioclase porphyroblasts with deformation twins, and talc-rich bands similar to the mafic schists. Dynamically recrystallized Caucete Group quartzitic mylonites with sparse muscovite and biotite lie below the thrust. The most abundant units in the sequence are strongly foliated, fine-grained massive amphibolitic mylonites, associated with lesser amounts of

170

G. I. VUJOVICH & S. M. KAY

500 100 50

Pie de Palo

a) 500 ~,~.,,.

~

9 PP25

10 5

Pie de Palo Complex b) Gneisses

l O C ~ _

10

',,Y/

o PP63 9 PP71 9 PP60

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9 PP25

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o PP63 9 PP71 . . . . 9 PP60

0.1

1

500

Quebrada del Gato Gneisses

loo[

c) 500 lOO~

Quebrada del Gato Gneisses PP99

50 ~

La Ce

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Sm Eu

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10

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d)

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Ta

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Nd

Sm

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Fig. 6. Chondrite-normalized REE and N-MORB normalized extended trace element patterns for Pie de Palo Complex gneisses. Protoliths are interpreted as arc-derived volcaniclastic sediments and intermediate to silicic arc magmas. (a) and (b) Gneisses from the Quebrada de Guayaupa/Lima interpreted as volcaniclastic sediments (PP60, 63, 71, and 93) and from the Nikizanga block interpreted as silicic lavas (PP25). (e) and (d) Quebrada del Gato gneisses interrupted as arc-related granitoids. PP98 is considered to be equilibrated with a plagioclase-poor, garnet-bearing residual mineralogy, taken as evidence for intrusion after crustal thickening caused by accretion of the Pie de Palo Complex to a larger crustal block. Further discussion in text. Data in Table 4. Normalization as in Fig. 3.

medium to coarser grained amphibolite. Under the microscope, these amphibolites are principally composed of prisms of actinolitic amphibole and subhedral to euhedral zoisite and epidote replacing plagioclase. Plagioclase is best preserved in the coarser varieties. Accessory minerals are titanite and sparse Fe-Ti oxides. Secondary quartz occurs in fine veins and recrystallized aggregates. On a macroscopic scale, alternating coarser and finer grained textures could be inherited from the effects of chilling in a sheeted dyke complex (E. M. Moores, pers. comm. 1995) consistent with the protolith of these amphibolites being mafic flows and dykes. Associated with the amphibolites are metagabbro/metadiorite bodies with mylonitized nematoblastic to granoblastic textures. They are primarily composed of tremolitic to actinolitic amphibole surrounding relict grains of clinopyroxene, and polysynthetically-twinned plagioclase that is almost totally replaced by epidote and sericite. Accessory minerals are titanite and tourmaline. Fine-grained, opaque-rich amphibole-talc schists with a prominent foliation marked by

tight microscopic scale folds and shear zones occur just above the fault. These schists are characterised by Fe-rich opaque grains that make up to 30-35% of the mode. The opaque minerals are associated with light green amphibole interspersed with epidote, and prismatic and laminar grains of talc. Large subhedral to euhedral grains of apatite, scarce subrounded grains of plagioclase with poorly developed twins, and lenses of recrystallized quartz and feldspar are also present. Tourmaline occurs as an accessory mineral. Chemically, the amphibolite samples (QQ1, QQ2, PR3, PR4a in Table 2 and Fig. 4a-b) are characterized by distinctively N-MORB like light-depleted REE patterns and Ta, Th, and Hf concentrations. In the H f - T a - T h plot in Fig. 7, they plot in the part of the N-MORB field occupied by Lau Basin back-arc spreading centre basalts. Complete major element analyses are not available, but FeO (10-12%) and Na20 (2.5-2.6%) concentrations are similar to those of slightly to moderately fractionated N-MORB basalt. Cr (80-310ppm) and Ni (45-80ppm) concentrations are consistent with preferential

GRENVILLE OCEANIC TERRANE IN PIE DE PALO Hf/3 inated luitas/Quemado nphibolite lafic schist Q. Guayaupa 9 arnphibilite central & eastern region amphibolite ~I, acidic gneiss~

Calc-

Th

o,25

0.50

0.75

Ta

Fig. 7. Hf/3-Th-Ta plot of Wood et al. (1979) showing distribution of Pie de Palo Complex mafic samples with melt-dominated trace element abundances relative to fields (in grey) for western Pacific Lau Basin region arc/back-arc samples from Hawkins (1995a, b). Also shown for reference are Pie de Palo Complex intermediate and silicic gneisses which all have arc characteristics. Points labelled with sample numbers are mentioned in text. See discussion in text. fractionation of olivine relative to clinopyroxene as expected under the low pressure crystallization conditions typical of N-MORB. Ba and Sr concentrations, are higher than expected for N-MORB, but are close to those of N-MORBlike back-arc lavas in the Lau Basin region (e.g., Hawkins 1995a, b). Cs and U, and to a lesser extent Ba and Sr, concentrations, have almost certainly been affected by secondary processes associated with metamorphism and deformation. An analyses of one of the white veins cutting the amphibolite unit is given in Table 2 (PR4c). This vein is speculatively interpreted as the highly altered, mylonitized remnants of a plagiogranite. Distinctive chemical features expected in a plagiogranite are high Na20 (9%) and low Sr concentrations, a flat REE pattern (La/Yb ratio < 5), an extreme negative Eu anomaly (Eu/Eu* < 0.1), and very high concentrations of Ta, Th, U, and Cs. A problem with this interpretation is a very low REE level (La c. 2 ppm). The gabbros (QQ4, PR4d, PR5 in Table 2 and Fig. 4c-d) have typical cumulate-dominated trace element patterns with low REE concentration. Their large positive Eu anomalies (Eu/Eu*= 1.7-2.2) and high Sr concentrations (340-380 ppm) are consistent with the presence of cumulate plagioclase. The higher Cr (550 ppm) and Ni (156ppm), lower REE level, and higher La/Sm ratio of sample QQ4 compared to the

171

others suggests that this gabbro was more clinopyroxene- and olivine-rich. Some characteristics of the gabbros including low La/Th (9-11) and high La/Ta (>30) ratios are more arc-like than those of the amphibolites. However, these ratios can be higher in cumulates than in lavas because Ta and Th are more incompatible and remain in the melt, whereas La is more compatible and enters pyroxene. The Ta concentration in QQ4 (<0.02ppm) is at the detection limit of the analyses. The chemistry of the mafic amphibole-talc schists (QQ6, 7 & 8 and PR2 in Table 2 and Fig. 4e-f) reflects their complex textures and mineralogy. Their opaque-rich mineralogy is reflected in very high FeO concentrations (14-25%), which are almost certainly the product of secondary processes associated with fluid circulation. With regard to generally immobile elements, samples PR2 (14% FeO) and QQ7 (19% FeO) have high Cr concentrations (260 and 360 ppm), light REE depleted patterns, and trace element levels like the amphibolites, whereas QQ6 (19% FeO) and QQ8 (25% FeO) have lower Cr (c. 50ppm) concentrations, a factor of ten times higher REE levels, and mild light-REE enrichment. In the Hf/Ta/Th plot (Fig. 7), QQ7 plots closer to the arc field than the amphibolites, and PP6 and PP8 plot near the border with the E-MORB field. PR2 has an extremely high Ta concentration that cannot be related to a melt composition. Although extreme deformation and mineralogical changes make the exact protolith of these samples difficult to interpret, they were certainly mafic igneous rocks formed in an oceanic environment. It is possible that PP2 and QQ7 were basaltic lavas similar to those represented by the amphibolites. QQ6 and QQ8 have a more E-MORB-Iike character and could represent seamount volcanism. A strong possibility is that the high FeO concentrations of these mafic schists predate the major metamorphic and deformational events that affected the Pie de Palo Complex, and are related to hydrothermal activity associated with spreading ridges (e.g., Ishibashi & Urabe 1995). In this case, high Fe and Co (57-60 ppm) concentrations, the high Ta concentration in PP2, and high REE concentrations in QQ6 and QQ7 could represent thalassic ferromanganese crusts like those described by Maillet et al. (1995) from the New Hebrides. The negative Ce anomaly and the lack of an Eu anomaly in the REE pattern of the most Fe-rich sample, QQ8, are consistent with this explanation. The presence of sparse tourmaline in these mafic schists and some of the metagabbros supports a role for fluid circulation. Some of this circulation could

172

G. I. VUJOVICH & S. M. KAY

be primary, but additional overprinting by metamorphic fluids circulating along the Las Pirquitas fault seems likely. Cs, Ba, Sr, U and Na concentrations are likely to have been affected by all of these processes. Overall, the protolith of the Quebradas Las Pirquitas/E1 Quemado mafic sequence is best interpreted as the top of a disrupted ophiolite with the opaque-rich talc-schists being hydrothermally altered mafic units, the amphibolite units being a sheeted dyke and pillow lava sequence, the metagabbro and metadiorite units being the cumulate section, and the Na-rich white veins, perhaps being plagiogranite. Given the regional setting, the difficulties of preserving and obducting N-MORB from the open ocean, and the chemical similarities with units in modern south-west Pacific back-arc basins like the Lau Basin (e.g., Hawkins 1995a, b), the most likely setting for this sequences is a back-arc spreading centre.

Mafic rocks associated with Pie de Palo Complex gneisses and marbles Arc, back-arc and shoshonitic magmatic rocks from the Quebrada del Molle, Mogote Corralitos and Nikizanga Block. Mafic amphibolites are also present in the central and eastern parts of the Sierra de Pie de Palo where they occur as tabular, elongate bodies of variable dimensions intercalated with biotite-muscovite-oligoclasegarnet schists and gneisses. They are typically fine- to medium-grained rocks with granoblastic to nematoblastic textures. The principal minerals are hornblende, epidote, plagioclase and quartz; accessory minerals are apatite, titanite, and zircon. Biotite is present in some cases. Analyses of samples from the Quebrada del Molle, in the Mogote Corralitos area, and in the extreme south-eastern part of the range are shown in Table 3 and plotted in Figs 5 and 7. As discussed by Ramos et al. (1993) and Vujovich & Kay (1996), the protoliths of most of these amphibolites can be interpreted as originating in an arc/back-arc environment. Starting in the west, Quebrada del Molle amphibolites PP13 and PP12, are chemically like arc lavas. The signature of PP13 is that of a typical high-A1 basalt from a modern oceanic arc, such as the Aleutians (e.g., Kay & Kay 1994). Distinctive features include high A1203 (17.5%) and low TiO2 (0.7%) concentrations, a relatively flat REE pattern (La/Yb = 2.2), essentially no Eu anomaly (Eu/Eu* = 0.98), and normalized concentrations of Ta that are low and Th, Ba and Sr that are high relative to La (Fig. 5a). The

signature of PP12 is that of an arc basaltic andesite. Lower REE, U and Th concentrations, and a higher La/Ta ratio (68) than for PP13 indicate a different parent magma. The analyses of amphibolite PP24 (52% SiO2) from the Nikizanga block in the SE part of the Sierra de Pie de Palo suggests an evolved arc or back-arc tholeiitic lava protolith (FeO/MgO = 1.6, Cr and Ni both <60 ppm; La/Ta = 30). Amphibolite PP3 from the Mogote Corralitos region can be interpreted as an arc-related magmatic rock, probably a cumulate. Distinctive arc features include a high La/Ta ratio (45), a moderately concave-upward REE pattern and high Sc concentration (51ppm) suggesting cumulate amphibole, and high Ti and Fe concentrations reflecting FeTi oxide phases. High Cr (247 ppm) and Sc concentrations suggest clinopyroxene. Low U, Th, and REE concentrations are consistent with a cumulate origin. Quebrada del Molle amphibolite PP10 (53% SiO2) from the central part of the range is distinctive in having high K20 (3.4%), light REE (244ppm Ce) and incompatible element concentrations (e.g., 1428ppm Ba) like those typical of shoshonitic lavas (Fig. 5a-b). A moderate Sr concentration (360ppm) associated with a relatively large negative Eu anomaly (Eu/Eu* = 0.72) and low Cr, Ni and Sr concentrations suggests that plagioclase and mafic minerals were fractionated from the primary melt. Steep HREE (Sm/Yb--3.7) and moderately steep overall REE (La/Yb = 13) slopes are consistent with small degrees of melt of a garnetbearing mantle source. Shoshonites like this are generally interpreted as melts of phlogopite garnet-bearing mantle enriched by subduction processes above the deeper parts of a subduction zone.

Diverse mafic magmas from the Quebrada Guayaupa-Lima. Two classes of mafic amphibolite occur in the Quebrada de GuayaupaLima. The first group consists of small, foliated, medium-grained, upper greenschist facies bodies associated with muscovite-biotite-plagioclasegarnet schists. Their principal mineralogy is clinopyroxene, hornblende, plagioclase, epidote and quartz; an accessory mineral is titanite. The second group occurs as lenses and bodies of fine to medium-grained amphibolite tectonically interleaved with marbles and micafeldspar-garnet gneiss. They have granoblastic to nematoblastic textures, and amphibolitefacies mineral assemblages containing hornblende, plagioclase, zoisite-clinozoisite, quartz, Fe-Ti oxides and apatite. Typical analyses are given in Table 4 and plotted in Fig. 5c-d.

GRENVILLE OCEANIC TERRANE IN PIE DE PALO Amphibolite PP62 in the first group, and PP73 in the second group, can be interpreted as magmatic rocks. PP62 (11.5% FeO and 3% Na20), a clinopyroxene-amphibole gabbro in terms of texture and mineralogy, shows segregation banding and occurs in a sequence of micaceous schists. PP73 (9.5% FeO; 2.9% Na20) is interfolded with a relatively pure marble. U and Sr spikes in its MORB-normalized pattern (Fig. 5d) probably reflect reaction with marble. The REE patterns of both samples are characterized by small negative Eu anomalies (Eu/Eu* = 0.86) and light REE depletion. Relative Ta/Th/Hf relations place them in the N-MORB field in Fig. 7, in accord with their non-arc like La/Ta ratios (15 to 17). Light REE depletion and very low Th concentrations argue for a depleted primary source. If these rocks were lavas, PP73 is more primitive than PP62 as shown by higher Cr (462 versus 82), and Ni (158 versus 71ppm) and lower REE, Ta, and Th concentrations. However, high Sc concentrations (46 to 49ppm) are consistent with cumulate pyroxene. Cumulate phases in PP73 could explain a low REE concentration relative to N-MORB (Fig. 5d), as well as a high Cr (462 ppm) concentration. Like the Las Pirquitas amphibolites, these samples are best interpreted as having formed in a back-arc rift setting. Two other Fe-rich samples, PP85 (14.7% FeO; 2.9% Na20) and PP92 (16% FeO; 2.2 Na20) could have cumulate or evolved magma protoliths (Cr and Ni < 65 ppm). Like PP73, both are complexly interfolded with relatively pure marble, and their Sr and U concentrations reflect various degrees of chemical interchange with marble. Features like their La/Ta (18-20) and Ta/Th ratios are consistent with a protolith in a back-arc spreading setting associated with samples PP62 and PP73. Alternatively, their protolith could be very proximal sediments. Sample PP79 from the Quebrada del Lima is an amphibole-clinopyroxene~lagioclasequartz gneiss. The REE pattern is characterized by a nearly flat HREE slope at <10 times chondritic abundance, a moderately positive LREE slope, and a negative Eu anomaly (Eu/Eu*= 0.80). Normalized low Ta and high Th relative to La concentrations, and a magmatic Th/U ratio (4.5) are consistent with a source in an arc environment. A proximal sedimentary protolith is favoured by siliciclastic-like Na20 (1.6%) and Sr (133ppm) concentrations, but not by a relatively low REE concentration. A secondarily-modified, relative primitive lava or clinopyroxene-olivine bearing, plagioclase-poor cumulate protolith with intercumulus liquid seems more likely.

173

Intermediate to acidic Pie de Palo Complex schists and gneisses Central and eastern schists/gneisses: greywackes and silicic magmas. Schists and gneisses are the dominant rock type in the central and eastern parts of the Sierra de Pie de Palo. The foliation in these typically fine- to medium-grained, light to dark grey rocks is expressed by alternating feldspar and mica-rich bands, some or which contain almandine garnet up to 0.5cm across. Textures vary from granoblastic to lepidioblastic. Principal minerals are quartz, oligoclase, biotite, and sometimes muscovite. Epidote and garnet also occur. Accessory minerals are apatite, titanite, and zircon. NE- and NWoriented, east-dipping mylonitic bands, varying from several centimetres to tens of metres in width, occur across the range. In some areas, the gneisses are proto-mylonites characterized by sub-rounded porphyroclasts of plagioclase, perthitic microcline, and garnet in a dynamically recrystallized matrix of quartz and biotite. Some porphyroclasts have asymmetric pressure trails of recrystallized quartz (typical 1-c mylonites). Four samples were analysed from along the length of the Quebrada de Guayaupa-Lima. From west to east, they include: PP60 (3.0% FeO, 3.4% Na20), a proto-mylonitic augen gneiss with a well-developed stretching lineation; PP63 (6.8% FeO, 3% Na20), a mylonitic garnetbiotite schist from a complexly folded region in a side canyon near Quebrada Lagarcito; PP71 (4.0% FeO; 2.9% Na20), a gneiss with an incipient migmatitic texture indicated by pink segregations and ptygmatic folding; and PP90 (4.6% FeO; 2.7% Na20), a migmatitic gneiss from the eastern side of the Sierra de Pie de Palo. Although a magmatic protolith cannot be ruled out, a siliciclastic origin seems probable for most, based on evidence for layering in the field and chemical comparisons with Phanerozoic siliciclastic rocks (Taylor & McLennan 1985). To a first order, the chemical characteristics of the analysed samples are similar (Table 4, Fig. 6a-b). Arc-like La/Ta (28-54) and La/Th ratios are consistent with formation in an arc environment. Siliciclastic-like features include nearly flat HREE and relatively steep LREE patterns, pronounced negative Eu anomalies, low Sr (180-206ppm) and moderate Na20 (2.73.4%) concentrations, low Cr (9-42 ppm) and Ni (10-53 ppm) and high Th and U concentrations, and typical Th/U ratios (4.0-5.4). In detail, the chemistry of PP60 is the least like a typical siliciclastic rock as it has the steepest REE pattern ( L a / Y b = 14.3), the most pronounced negative Eu anomaly (Eu/Eu* = 0.40, siliciclastic

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G. I. VUJOVICH & S. M. KAY

typically=0.57-0.63), the highest Na20 (3.4%) concentration, and the lowest FeO (3.0%) and Cr (9ppm, average shale >50ppm) concentrations. Its protolith could be an arc dacite, a granodiorite, or a sediment derived from such rocks. Other gneisses could have magmatic protoliths. In particular, silicic mylonite PP25 (66% SiO2) from the Nikizanga block was interpreted as a dacitic magma by Ramos et al. (1993) based on a typical arc rhyodacite trace element pattern. An argument supporting a magmatic origin is a high Na20 (4.2%) concentration. Zircons used in dating a similar sample were euhedral, doubly terminated crystals (McDonough et al. 1993). Quebrada del Gato arc rocks." arc magmas erupted through a thickened crust. Intermediate to silicic

gneisses in the Quebrada del Gato are distinct from those in the central and eastern Pie de Palo Complex. The protoliths of these gneisses were intrusive into the Quebrada del Gato mafic/ultramafic sequence as seen at the low cliff at the lower end of the quebrada where medium-grained, light grey gneisses like sample PP98 contain xenoliths of metamorphosed mafic rocks. These gneisses are predominantly composed of tremolitic to actinolitic hornblende, and plagioclase altered to aggregates of epidote (zoisite-clinozoisite). Grey, medium-grained, biotite-bearing tonalitic-granodiorite gneisses, like samples PP99 and PP108, which crop out further up the quebrada also contain mafic xenoliths. Among the chemical characteristics that distinguish the Quebrada del Gato gneisses from those studied from the central and eastern Pie de Palo Complex are their higher Na20 (5.1-5.9%) and Sr (420-550ppm) concentrations, much higher La/Ta ratios (80-115), generally steeper light REE (La/Sm = 5-8) and overall REE (La/ Y b = 12-44) patterns, lower heavy REE abundances, and lack of negative Eu anomalies (compare Fig. 6a-b with c-d). High Na20 and Sr concentrations are consistent with their protoliths being igneous rocks. Several chemical characteristics of samples PP98, PP99, and PP108 (Table 4; Fig. 6c-d), like high Ba/La (20-27) and very high La/Ta (>80) ratios, are consistent with them all having arc granodioritic protoliths. However, important differences occur among them. In particular, a positive Eu anomaly (Eu/Eu* = 1.24), and high AlzO3 (19%) and Sr (542ppm) concentrations point to PP108 being plagioclase-rich. Abundant plagioclase and a relatively flat REE pattern (La/Yb= 12) indicate low pressure differentiation consistent with evolution in an arc on

oceanic or attenuated continental crust. In contrast, the other samples have lower overall REE concentrations, steeper REE patterns, and virtually no Eu anomalies. The pattern of PP99 (La/Yb = 21), and particularly that of PP98 (La/Yb=44), is consistent with equilibration with a high pressure mineral residua that was plagioclase-poor, as indicated by the lack of an Eu anomaly, and garnet-bearing, as indicated by a steep HREE pattern and low HREE abundances. Such a situation is expected if mafic arcderived magmas equilibrate with high-pressure garnet-granulite or eclogite facies units in a thickened crust. Relatively high Th, U, and Ba concentrations support crustal contamination. Samples PP98 and PP99 are difficult to ascribe to the same arc/back-arc setting on thin crust as the other Pie de Palo Complex units. A solution is for these magmas to have been emplaced in an arc regime superimposed on older units after crustal shortening and thickening by compressional deformation. Such a deformational event could be associated with collision and accretion of a terrane containing the Pie de Palo Complex to a larger continental block. Detailed structural analyses and geochronology are needed to evaluate the geometry and timing of such a possible collision

Discussion: a Mesoproterozoic oceanic arc/back-arc setting for the Pie de Palo Complex All of the protolith types assigned to the Pie de Palo Complex, except that of the Quebrada de! Gato silicic gneisses, can all be found among the arc and marginal basin units of the western Pacific region. This correspondence lends credence to the proposal that the Pie de Palo Complex formed in an oceanic arc/back-arc setting in Mesoproterozoic times. To explore this proposal, the units of the Pie de Palo Complex are compared with those of the Lau Basin region, whose characteristics and evolution have been reviewed by Hawkins (1995a, b). A cross-section across the Lau Basin region after Hawkins (1995b) is shown in Fig. 8 with the proposed Pie de Palo Complex analogues at the base. An important factor in interpreting the protolith of the Pie de Palo Complex is the time required for emplacement of the major units. Existing age constraints indicate that most of the dated Pie de Palo Complex units were emplaced between 1090 and 1060Ma. Including the single zircon fraction 2~176 age of 1118 Ma for the Cerro Barboza schist, the time span is around 60 million years. Interestingly,

GRENVILLE OCEANIC TERRANE IN PIE DE PALO the arc and back-arc development recorded in the Lau Basin-Tonga region spans more than 40Ma (Hawkins 1995b), which is roughly the same order of magnitude. In detail, Hawkins (1995a, b) suggests that the modern Lau-Tonga supra-subduction system began forming in the Early to Mid-Miocene. The Eocene units in the fore-arc formed in older convergence systems that were carried along by subsequent back-arc rifting. These earlier subduction systems were probably located well to the west; their polarity is uncertain. The protolith of the Pie de Palo Complex probably formed in a similarly complex system. A critical factor in comparing the Pie de Palo Complex with the Lau Basin region is the chemistry of the mafic magmas. To this end, some key chemical characteristics of Lau Basin region magmas are compared with those of Pie de Palo Complex samples in the Hf/Th/Ta discrimination diagram of Wood et al. (1979) in Fig. 7. Importantly, amphibolites with N-MORB like light REE depleted patterns from the Quebradas Las Pirquitas/E1 Quemado and Guayaupa plot in the same field as light REE depleted lavas from back-arc ridges in the Lau region (kau Spreading Centre in Fig. 8). The Quebrada Las Pirquitas Fe-rich mafic schists lie on the margin of this field. These samples and the Quebrada Guayaupa amphibolites that fall in the E-MORB field have compositions like Peggy Ridge and Rochambeau Bank magmas in the northern Lau Basin (not shown in Fig. 7; see Hawkins 1995a). Equally important, the Valdivia cumulates (V44, V45) with melt-dominated trace element patterns fall in the field of the Tofua field, which is representative of Lau region arc lavas. Other cumulates from Cerro Valdivia (V43) and Quebrada del Gato (P100) with melt-dominated trace element tendencies, and amphibolites from Quebrada Molle (P10 and P12) and Mogote Corralitos (P3) plot in the arc region near the Tofua field. The only mafic Pie de Palo Complex samples without Lau region equivalents are Quebrada Guayaupa amphibolite P79, which could be a greywacke, and the shoshonitic-like Quebrada del Molle sample P13. Shoshonitic lavas are found in other parts of the western Pacific region (e.g., seamounts in the Marianas trough, see Stern et al. 1988). The Quebrada del Gato and Cerros Valdivia/ Barboza arc cumulates resemble units found on the Lau Ridge. Uplift of such a ridge is consistent with erosion of lavas from the top, exposing the arc cumulate section. The other major Pie de Palo Complex protolith types also occur in the Lau region. The Quebrada Las Pirquitas Fe-rich mafic schists are

175

consistent with hydrothermal venting activity which is common in back-arc rifts. The Pie de Palo marbles are explicable as stromatolitic reefs equivalent to coral reefs under arc-derived volcaniclastic rocks in the Lau fore-arc or on the Lau ridge (Fig. 8). The greywackes are explicable as volcaniclastic sediments shed from an arc split by back-arc rifting (see Clift 1995). Melting of older arc crust is thought to be largely responsible for the dacitic and rhyolitic magmas underlying the volcaniclastic sequence in the fore-arc of the Tofua arc (Fig. 8; Hawkins 1995b). Sample QQ25 whose protolith is interpreted as a dacite could be equivalent to these units. As expected for crustal melts in an arc region and immature greywackes eroded from an arc source, the Pie de Palo gneisses have arc trace element signatures. The only samples that do not fit the Lau region analogy are the Quebrada del Gato silicic gneisses with steep REE patterns. Their trace element characteristics could be explained if they were younger plutonic rocks emplaced in a crust thickened by collision of the Pie de Palo arc/back-arc basin complex with a larger continental block. The above comparison with the Lau Basin region implies that the Mesoproterozoic subducting slab dipped westward relative to the modern Sierra de Pie de Palo as shown in Fig. 8, but the possibility that the slab dipped eastward cannot be ruled out. In the latter case, the Quebrada del Gato/Cerros Valdivia and Barboza units would be in the position of the Tofua arc in the Lau region, and the central and eastern gneisses and amphibolites in the back-arc. Westverging structures in the Sierra de Pie de Palo region could be taken as evidence for this orientation, since west-vergence is the general case over an east-dipping subduction zone. The major structures in the Sierra de Pie de Palo could indicate a Mesoproterozoic collision between a Pie de Palo arc/back-arc complex over an east-dipping subduction zone and a terrane to the west. A problem is that the age of the major structure relative to the arc/back-arc sequence is unknown.

Implications for Gondwana-Laurentian Terrane interactions from the Pie de Palo Complex An important regional tectonic question is the relation between the Pie de Palo Complex and the kaurentian-derived Precordillera basement to the west (Fig. 1). A major Mesoproterozoic tectono-magmatic event is reflected by Grenville-

176

G. I. VUJOVICH & S. M. KAY

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Mesoproterozoic Pie de Palo Complex Equivalents Fig. 8. Cross-section of the Lau Basin region modified from Hawkins (1995b) showing the culmination of >40 million years of arc/back-arc history. All of the protolith types seen in the Mesoproterozoic Pie de Palo Complex are present. The inset at the bottom shows the correspondence of the various parts of the Pie de Palo Complex to the Lau Basin region given a west-dipping subduction relative to the modern Sierra de Pie de Palo. The possibility that the subduction zone dipped to the east also exists. In that case, the Quebrada del Gato, Barboza and Valdivia arc rocks, shown in the position of the extinct arc, could be switched with the Central and Eastern Pie de Palo units. age (c. l l00-1050Ma) crystallization and metamorphic ages in both ranges. A key question is whether or not the Pie de Palo Complex was transported from Laurentia to Gondwana along with the Precordillera Terrane. Key problems in understanding the relation between the Precordillera and Pie de Palo Complex are the age of the ductile thrusting that puts the Pie de Palo Complex over the low grade siliciclastic sediments and limestones of the Caucete Group, and the age and origin of the Caucete Group. The last ductile motion on the thrust is known to be c. 395 Ma (see Ramos et al. this volume), but the history of earlier motions is unknown. Available constraints allow the age of the Caucete Group to be anywhere between Grenvillian and early

Palaeozoic (see Ramos et al. 1996). A further complication is that the region between the Pie de Palo Complex and the Precordillera is mostly covered by Miocene to Recent sediments. To the west, the eastern part of the Precordillera consists of the Villicum and Zonda ranges (Fig. 1) which are largely composed of Cambrian to Ordovician sedimentary rocks. The Lower Cambrian units in these ranges have palaeontological and stratigraphical affinities to Laurentia (e.g., Astini et al. 1995). They are overlain by platform carbonates associated with a passive margin that endured until the lower Ordovician, and then by Ordovician to Devonian sediments that accumulated in a back-arc basin (e.g., Astini et al. 1995, 1996). Palaeontological evidence

GRENVILLE OCEANIC TERRANE IN PIE DE PALO indicates that the eastern Precordillera block was part of Gondwana by the mid-Ordovician (e.g., summary in Dalziel et al. 1996). Evidence for an Ordovician Sierra de Pie de PaloPrecordillera suture is not apparent from radiometric dating in the Caucete Group. It is possible that the strong regional Devonian thermal event recorded by 4~176 white mica ages of c. 395 Ma (Ramos et al. this volume) overprinted an Ordovician event. This Devonian event is generally associated with collision of the Chilenia terrane with the Precordillera to the west (Ramos et al. 1986, 1996). A scenario, which is not mutually exclusive with overprinting of an Ordovician event, is that the Precordillera and Sierra de Pie de Palo blocks were together in the Cuyania terrane (see Ramos et al. 1996) by Early Palaeozoic time, and that the Gondwana-Laurentian suture is on the eastern side of the Sierra de Pie de Palo (e.g., Dalla Salda et al. 1992; Astini et al. 1995; Kay et al. 1996; see Fig. 1). In the latter scenario, Grenville-age deformation and metamorphism could mark suturing of the Pie de Palo Complex with the Precordillera block. The collision would have had to occur on the eastern margin of Laurentia as the Precordillera was attached to Laurentia until latest Precambrian time (e.g., Ramos et al. 1993; Astini et al. 1995, 1996; Dalziel et al. 1996). Subsequently, the Cuyania terrane, composed of the Precordillera and the Sierra de Pie de Palo block, could have rifted from Laurentia to be sutured to Gondwana in the Early Palaeozoic. The most likely site of detachment of the Precordillera is the Ouachita region, east of the Llano uplift in Texas in the United States (e.g., Astini et al. 1995; Kay et al. 1996; Dalziel et al. 1996). Such a scenario is attractive as an explanation for the juvenile arc/back-arc protoliths with Grenville metamorphic overprints that occur in both the Precordillera (Kay et al. 1996) and the Sierra de Pie de Palo, but that are largely missing in the Grenville basement of eastern North America. This hypothesis needs to be tested with constraints on the age of the west-verging structures of the Pie de Palo Complex and the Caucete Group, and the location of the Lower Palaeozoic GondwanaLaurentian suture.

Conclusions Conclusions from this field, petrographical and geochemical study of the units of the Pie de Palo Complex are as follows. (1) Patterns of immobile and some mobile trace elements are useful in establishing the protoliths of the heavily deformed, greenschist

177

to amphibolite facies, metamorphic rocks in the Pie de Palo Complex. (2) The Pie de Palo Complex represents a Mesoproterozoic Grenville-age oceanic arc/ back-arc complex similar to that formed in the western Pacific Lau Basin region. Existing radiometric constraints suggest that the protolith of the Pie de Palo Complex, like Eocene to Recent Lau Basin units, was emplaced in a time frame on the order of 40 to 60 million years. (3) The juvenile ultramafic and mafic magmatic rocks in the Pie de Palo Complex, along with those in the Precordillera basement to the west, are among the most primitive Mesoproterozoic magmatic rocks known on Earth. (4) The mafic sequence in the Quebradas Las Pirquitas/E1 Quemado region represents the basaltic and plutonic section of a back-arc spreading ridge ophiolite in which lavas have N-MORB like light REE depleted trace element patterns and ratios involving La, Th, Ta, and Hf. Mafic Fe-rich schists in the sequence could reflect hydrothermal processes at the spreading centre. Evidence for a back-arc origin comes from association with the other lithological elements expected in a Lau Basin type arc/back-arc complex. (5) The mafic-ultramafic sequences in the Quebrada del Gato and at Cerros Barboza and Valdivia are largely composed of the cumulate sections of an island arc ophiolite. An arc setting is indicated by evidence for primary amphibole and chromite, and characteristic Ta, Th, Hf and REE ratios in melt-dominated trace element patterns. The lack of obvious lava counterparts can be explained by uplift of the extinct arc after it was rifted by back-arc spreading. Uplift could have been accentuated by collision with a terrane to the west. (6) Amphibolite bodies associated with gneisses in the central and eastern parts of the Sierra de Pie de Palo have trace element signatures indicating that a mixed arc and back-arc origin. Most are best interpreted as lavas and cumulates. One has the characteristics of an island arc shoshonite. (7) Intermediate to silicic Pie de Palo Complex gneisses probably largely originated as arcderived sedimentary units (greywacke). Melting of arc crust at the time of arc splitting can account for gneisses with silicic magmatic affinities. Marbles are best interpreted as stromatolitic reefs. (8) Intermediate to silicic gneisses with calcalkaline arc signatures in the Quebrada del Gato are intrusive into the mafic-ultramafic sequence. Trace element evidence for a high pressure mineral residue in some is consistent

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G. I, V U J O V I C H & S. M. K A Y

with e m p l a c e m e n t in an arc crust thickened by a m a j o r c o m p r e s s i o n a l d e f o r m a t i o n . This event c o u l d reflect s u t u r i n g o f the Pie de Palo C o m plex to a n o t h e r terrane. (9) Evidence for juvenile a r c / b a c k - a r c p r o t o liths with Grenville crystallization a n d metam o r p h i c ages in b o t h the Pie de Palo C o m p l e x a n d the Precordillera b a s e m e n t are consistent with a M e s o p r o t e r o z o i c collision o f these terranes a l o n g the L a u r e n t i a n margin. Both w o u l d have been a t t a c h e d to the L a u r e n t i a n m a r g i n o f the m o d e r n s o u t h e a s t e r n U S in the Late Prec a m b r i a n w h e n they were rifted f r o m L a u r e n t i a as the C u y a n i a Terrane. This implies that the L o w e r Palaeozoic G o n d w a n a - L a u r e n t i a n suture lies east o f the Sierra de Pie de Palo as p r o p o s e d in a n u m b e r of p a p e r s (e.g., Dalla Salda et al. 1992; Astini et al. 1995; K a y et al. 1996; R a m o s et al. 1996). Partial funding for this project was provided by the following g r a n t s - US National Science Foundation EAR92-05042, UBACYT Ex-132 and CONICET PID 4162. We thank R.W. Kay, V. A. Ramos, and Eldridge Moores for discussion, J. Abbruzzi and M. L. Gorring for help with analytical work, and M. Godeas, N. Pezzutti and L. Villar for help with petrography. R. Pankhurst, T. Brewer, and especially J. Hawkins provided comments which improved the manuscript. The Argentine Servicio Geoldgico Nacional and Fundaci6n Antorchas supported the field work. This paper is a contribution to IGCP 345 and IGCP 376 'Laurentian-Gondwana Connections before Pangea'.

Appendix: analytical methods Whole-rock samples were pulverized in an aluminium oxide ceramic shatter-box. Electron microprobe analysis for major elements, and Instrumental Neutron Activation Analyses (INAA) for trace elements were carried at Cornell University. Glasses were prepared for major element analysis from c. 1 g of powder fused in a molybdenum strip furnace or fused with a gi metaborate flux in carbon crucibles. Analyses were performed on a three-spectrometer JEOL-733 Superprobe in wavelength dispersive spectrometric (WDS) mode and are based on averages of 4 to 6, 30-ram spot analyses. Operating conditions were 15 kV accelerating voltage and 15nA beam current. Standards were natural mineral and basaltic glasses. Data were reduced on a Tracor-Northern system using Bence-Albee matrix corrections. Typical two-sigma precision and accuracy is 1-5% for major elements (e.g., >1 wt%) and c. 5 10% for minor elements (e.g., <1 wt%) based on replicate analysis of glass standards. Trace element analysis by INAA were carried out on 0.5 g of sample powder packed in ultrapure Suprasil quartz tubing. Samples were packed along with three internal standards and irradiated in a TRIGA reactor (Ward Laboratory, Cornell University) at a power level of

c. 400 kW for 3 to 4 hours. Samples and standards were counted for 4 to 10 hours on an Ortec Intrinsic Ge detector: 6 and c. 40 days after irradiation. Precision and accuracy (two sigma) based on replicate standard analysis are 2- 5% for most elements; c. 8% for U, Sr, and Nd. Actual precision depends on counting time relative to concentration of element. Further information on techniques and standards used for INAA analyses is in Kay et al. (1987).

References ABBRUZZI, J. M., KAY, S. M. & BICKFORD, S. M. 1993. Implications for the nature of the Precordilleran basement from the geochemistry and age of Precambrian xenoliths in Miocene volcanic rocks, San Juan province. XII Congreso Geol6gico Argentino y II Congreso de Exploraci6n de Hidrocarbm'os, Mendoza, Actas, III, 331-339. ASTINI, R. A., BENEDETTO,J. L. ~9.,VACCARI, N. E. 1995. The Early Paleozoic evolution of the Argentine Precordillera as a Laurentian rifted, drifted and collided terrane: A geodynamic model. Geological SocieO' of America Bulletin, 107, 235 273. , RAMOS, V. A., BENEDETTO,J. L., VACCAR[,N. E. & CANAS, F. L. ] 996. La Precordillera: un terreno ex6tico a Gondwana. XIII Congreso Geolbgico Argentino y Ill Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, V, 293 323. CASTRO DE MACHUCA, B. 198l. Ghnesis de l a m i n a de talco "Don Le6n", sierra de Pie de Palo, provincia de San Juan. VIII Congreso Geol6gico Argentino, San Luis, Actas, IV, 535-555, -& CONTE-GRAND, A. 1994. G4nesis de las antibolitas de las Sierras Pampeanas occidentales, provincia de San Juan, Argentina, 7~ Congreso Geol@ico Chih, no, Concepci6n, Actas, II, 989 994. --, MEISSL, E. & CONTE-GRAND, A. 1995. Las serpentinitas de la sierra de Pie de Palo y su mineralizaci6n, provincia de San Juan, Argentina. V Congreso National de Geologia Econbmica, San Juan, Actas, 1, 1-11. CLIFT, P. D. 1995. Volcaniclastic sedimentation and volcanism during the rifting of western Pacific back arc basins. In: TAYLOR, B. & NATLAND, J. (eds) Active Margins and Marginal Basins of the Western PacO%. American Geophysical Union, Geophysical Monographs, 88, 67-96. DALLA SALDA, L. & VARELA, R. 1982. La estructura del basamento del tercio sur de la Sierra de Pie de Palo, provincia de San Juan. V Congreso Latinoanwricano de Geologia, Buenos Aires, Actas, 5, 451-468. -& VAaELA, R. 1984. E1 metamorfismo en el tercio sur de la sierra de Pie de Palo, San Juan. Revista de la Asociaci6n Geolbgica Argentina, Buenos Aires, 39, 68-93. - , DALZ]FL,I. W. D., CINGOLANI,C. A. & VARELA,R. 1992. Did the Taconic Appalachians continue into southern South America? Geology, 20, 1059-1062.

G R E N V I L L E O C E A N I C T E R R A N E IN PIE D E P A L O DALZIEL, I. W. D., DALLASALDA,L. H., CINGOLANI,C. & PALMER, A. R. 1996. The Argentine Precordillera: A Laurentian terrane. GSA Today, 6, 16-18. HAWKINS, J. W. 1995a. Evolution of the Lau Basin Insights from ODP Leg 135. In: TAYLOR, B. & NATLAND, J. (eds) Active Margins and Marginal Basins of the Western Pacific. American Geophysical Union, Geophysical Monographs, 88, 125-169. - - 1 9 9 5 b . The geology of the Lau Basin. In: TAYLOR, B. (ed.) Backarc Basins: Tectonics and May(matism. Plenum Press, New York, 63-129. ISHIBAStil, J.-I. & URABE, T. 1995. Hydrothermal activity related to arc-backarc magmatism in the western Pacific. In." TAYLOR, B. (ed.) Backarc Basins: Tectonics and Magmatism. Plenum Press, New York, 497 517. JORDAN, T. E. & ALLMENDINGER, R. W. 1986. The Sierras Pampeanas of Argentina: a modern analogue of Laramide deformation. American Journal of Science, 286, 737 764. KAY, S. M. & KAY, R. W. 1994. Aleutian magmatism in space and time. In: PLAFKER, G. & BERG, H. C. (eds) The Geology of Alaska, Decade of North American Geology (DNAG), Geological Society of America, G-I, 687-722. , MAKSAEV, V., MPODOZIS, C., Moscoso, R. & NASL C. 1987. Probing the evolving Andean lithosphere: middle to late Tertiary magmatic rocks in Chile over the modern zone of subhorizontal subduction (29-31.YS). Journal o/" Geophysical Research, 92, 6173-6189. , ORRELL, S. & ABBRUZZl,J. M. 1996. Zircon and whole rock Nd-Pb isotopic evidence for a Grenville age and Laurentian origin for the basement of the Precordillera terrane in Argentina. Journal of Geology, 104, 637-648. LEVERATTO, M. A. 1968. Geologia de la zona al oeste de U1Km-Zonda, horde oriental de la precordillera de San Juan, eruptividad subvolc'finica y estructura. Revista de ta Asociaci6n GeottJgica Argentina, 23, 129-157. LLANO, J. & ESCALANTE, A. 1979. Petrologia de las anfibolitas de la quebrada del Gato, sierra de Pie de Palo. 2 '~ Congreso Geol6gico Chileno, Arica, Actas, IlI, E39-E58. MAILLET, P, RUELLAN, E., GERARD, M., PERSON, A., BELLON,H., COTTON,J., JORON,J.-L., NAKADA,S. & PRICE, R. C. 1995. Tectonics, magmatism, and evolution of the New Hebrides backarc troughs (southwest Pacific). In: TAYLOR, B. (ed.) Backarc Basins." Tectonics and Magmatism. Plenum Press, New York, 170-230. MCDONOUGH, M. R., RAMOS, V. A. ISACHSEN, C. E., BOWRING, S. A. & VUJOWCH, G. I. 1993. Nuevas edades de circones del basamento de la sierra de Pie de Palo, Sierras Pampeanas Occidentales de San Juan: sus implicancias para los modelos del supercontinente proterozoico de Rodinia. XII Congreso Geol6gico Argentino y H Congreso de Exploraci6n de Hidrocarburos, Mendoza, Actas, IIl, 340 342. RAMOS, V. A. & VUJOWCH, G. 1998. Hoja Geol6gica San Juan, escala 1:250.000. Subsecretaria de

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Mineria, Direccidn Nacional del Servicio Geoldgico, Buenos Aires, in press. --, DALLMEYER, R. D. & VUJOVlCH, G. I. 1998. Time constraints in the Early Palaeozoic docking of the Precordillera, central Argentina. This l'ollllllC.

JORDAN, T. E., ALLMENDINGER, R. W., MPODOZ~S, C., KAY, S. M., CORTES, J. M. & PALMA, M. A. 1986. Paleozoic terranes of the central Argentine-Chilean Andes. Tectonics, B, 855 880. , VUJOVlCH, G. I., KAY, S. M. & MCDONOUGH, M. R. 1993. La orog~nesis de Grenville en las Sierras Pampeanas Occidentales: Pie de Palo un sistema de acreci6n de arcos islfindicos. Y l I Congreso Geol6gico Arg,entino y H Congreso de Exploraci(hl de Hidrocarburos, Mendoza, Actas, !11, 343 357. & DALLMEYER, D. 1996. Los klippes y ventanas tect6nicas de la estructura prefindica de la sierra de Pie de Palo (San Juan): Edad e implicancias tectdnicas. XIII Congreso Geol6gico Argentino v IlI Congreso de Exploracian de Hieh'ocarburos, Actas, V, 377-391. ROLLINSON, H. 1993. Using Geochemical Data. Longman Scientific and Technical, New York. SUN, S.-S. & MCDONOUGH, W. F. 1989. Chemical and isotopic systematics of oceanic basalts; implications for mantle composition and processes. In: SAUNDERS, A. D. & SORRY, M. J. (eds) Magmatism in the Ocean Basins. Geological Society, London, Special Publications, 42, 313-345. STERN, R. J., BLOOMER, S. H., LIN, P.-N., ITO, E. & MORRIS, J. 1988. Shoshonitic magmas in nascent arcs: new evidence from submarine volcanoes in the northern Marianas. Geology, 16, 426-430. TAYLOR, S. R, & MCLENNAN, S. M. 1985. The Continental Crust." its Composition and Evolution. Blackwell Scientific Publications, Oxford. VARELA, R. & DALLASALDA, L. 1993. Geocronologia Rb-Sr de metamorfitas y granitoides del extremo sur de la Sierra de Pie de Palo, San Juan. Revista de la Asociaci6n Geoldgica Argentina, 47, 271-276. VILLAR, L. 1985. Las rajas ultrab'asicas argentinas, tipos de ultram'aficas. Metalogenia. 4 ~ Congreso Geol6gico Chileno, Antofagasta, Actas, II, 610-633. VUJOVlCH, G. I. & KAY, S. M. 1996. Evidencias geoquimicas del origen y ambiente geol6gico de las rocas metam6rficas de composici6n m~ifica a intermedia de las Sierras Pampeanas Occidentales. XIII Congreso Geol6gico Argentino y III Congreso de Exploraci~Jn de Hidrocarburos, Buenos Aires, Actas, V, 273 291. & RAMOS, V. A. 1994. La faja de Angaco y su relaci6n con las Sierras Pampeanas Occidentales. 7 Congreso Geoh~gico Chileno, Concepcion, Actas, I, 215-219. WILSON, M. 1989. Igneous Petrogenesis. Unwin Hyman, London. WOOD, D. A., JORON, J. L. & TREVIL, M. 1979. A reappraisal of the use of trace elements to classify and discriminate between magma series erupted in different tectonic settings. Earth and Planetary Science Letters, 45, 326 336. -

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The Pampean Orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de C6rdoba C. W. R A P E L A 1, R. J. P A N K H U R S T 2, C. C A S Q U E T 3, E. B A L D O

4,

J. S A A V E D R A 5, C. G A L I N D O 3 & C. M. F A N N I N G 6

1 Centro de Investigaciones Geol6gicas, Universidad Nacional de La Plata, Calle 1 No. 644, 1900 La Plata, Argentina 2 British Antarctic Survey, N E R C Isotope Geosciences Laboratory, Keyworth, Nottingham NG12 5GG, UK 3 Departamento de Petrologia y Geoquimica, Universidad Complutense de Madrid, Madrid, Spain 4 Departamento de Geologia, Facultad de Ciencias Exactas, Fisicas y Naturales, Universidad Nacional de C6rdoba, Argentina 5 CSIC, Instituto de Agrobiologia y Recursos Naturales, Salamanca, Spain 6 Research School of Earth Sciences, The Australian National University, Canberra, Australia Abstract: A detailed study of the pre-Silurian geology of the Sierras de C6rdoba, Eastern Sierras Pampeanas, is used to define the sequence of magmatic and metamorphic events during the Pampean orogeny. This primarily involved Early to Mid-Cambrian subduction and terrane collision at the western margin of Gondwana during the amalgamation of the super-continent. Evidence for this is based principally on new information concerning (a) regional mapping and field relations, (b) analysis of the structures, deformational history and meta-morphic evolution and (c) geochronology and geochemistry of the igneous and metamorphic rocks. The main events recognized are (1) Late Proterozoic break-up of Rodinia (Nd model ages of 1500 + 200 Ma, inherited zircons 800-1400 Ma), (2) development of an Early Cam-brian passive margin sequence (Puncoviscana Formation and equivalents), (3) emplacement of metaluminous calc-alkaline granitoids (Gla, dated at 530 + 3 Ma) as a result of NE-directed subduction, (4) crustal thickening, ophiolite obduction, compression and high-grade metamorphism (M2: 8.6+0.8kbar, 8104- 50~ c. 525 Ma) related to collision, and culmina-ting in (5) isothermal uplift and widespread low-P anatexis (M3, 4.0 + 0.5 kbar, 715 + 15~ c. 520 Ma). The last event is responsible for the linked generation of highly peraluminous granites (Glb) and cordieritites. Subsequent emplacement into the accreted terrane of Ordovician trondhjemite-tonalites (500-470Ma) and dextral wrench shear are interpreted as inner cordilleran counterparts of the Famatinian arc, which developed to the west along the newly-formed proto-Andean margin.

The early geodynamic evolution of the Pacific margin of South America is still poorly constrained due to the scarcity of precise geochronological data for the metamorphic sequences and detailed petrological studies of related igneous suites. Reconstruction of P-T-t paths in cross-sections along the Sierras Pampeanas seems to be an unavoidable task if the history of the Gondwana foreland has to be disentangled from that of the supposed terranes that were accreted during the Pampean (Cambrian) and Famatinian (Ordovician) orogenies. With this main goal in mind, a series of isotopic, petrological and sedimentological studies in a geotraverse from the Eastern Sierras Pampeanas to the

Precordillera (PAMPRE Project) is being carried out by an interdisciplinary team of geologists mainly from Argentine, British and Spanish institutions. Unravelling the geological history of the polymetamorphic basement of the Sierras de C6rdoba (SC) at the eastern margin of the Sierras Pampeanas (Fig. 1) has been one of the main objectives of PAMPRE, results from which are presented in this paper. The sequence and characteristics of pre-Silurian metamorphic and magmatic events in the SC (Fig. 2) are inferred from (i) isotope dating of metamorphic and igneous rocks, using U - P b on abraded zircons, U-Pb SHRIMP analyses, and Rb-Sr whole-rock

RAPELA, C. W., PANKHURST, R. J., CASQUET, C. BALDO, E. SAAVEDRA,J. GALINDO, C. &; FANNING, C. M. 1998. The Pampean Orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de Cdrdoba. In: PANKHURST,R. J. & RAPELA,C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 181-217.

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Fig. 1. Geological map of the Sierras de Cdrdoba, modified after Kraemer et al. (1995). For regional context of the Sierras Pampeanas and Precordillera, see Pankhurst & Rapela (this volume).

and K-Ar determinations, (ii) thermo-barometric calculations based on mineral analyses, and (iii) the isotopic (Nd and Sr) and major and trace element geochemical signatures of the different magmatic suites. The geodynamic interpretation emerging from these data indicates that the proto-Andean margin at this latitude was the site of convergence and collision of continental blocks during Early and Mid-Cambrian times. After a period of quiescence during the Late Cambrian, the Sierras de Cdrdoba underwent brittle deformation and intrusion of scattered Early Ordovician plutons, representing an early stage of the main Famatinian arc located to the west, and emplaced through the accreted Cambrian terrane.

Geological overview of the Sierras de C6rdoba Lithology

The Sierras de Cordoba form one of the easternmost blocks of the Sierras Pampeanas geological province in central western Argentina (Fig. 1). As elsewhere in this province, present-day landforms result from basement uplift on reverse faults during an upper Tertiary stage of the Andean orogeny; a Quaternary plain extends for several hundred kilometres to the east. The SC basement is largely composed of polymetamorphic rocks, granitoid plutons emplaced at various levels, and subordinate basic and

THE PAMPEAN OROGENY OROGENY/ AGE (Ma)

Z <

360-370

.~435-447

DEFORMATION

METAMORPHISM

Contact metamorphism

G3: Post-orogenic, high-K batholiths and plutons Pegmatite III

D3: Upright folds, $3 foliation, dextral wrench shear

Low-grade mylonites

Pegmatite swarms (Pegmatite II) G2: High-Na trondhjemitic and granodioritic plutons

Extensional shear zones and uplift

M3: Isothermal decompression Low-P (2.5 - 4.0 kb) High-grade mylonites, Glb: Strongly migmatites and peraluminous granites, widespread anatexis Pegmatite 1

D2: syn-metamorphic, progressive, coaxial deformation ($2 regional foliation) Local reverse shearing

M2: High-T/medium-P (7- 9 kb) Gneiss and migmatite

519-525

Ophiolite obduction

Gla: Metaluminous, calc-alkaline granitoid sequences

527-533 DI' $1 relict foliation (unknown age) 540-570

MAGMATISM

D4: Shear zones

480-500

Z < IJJ Q_ 520-530 :E

183

MI: Medium-grade

Neoproterozoic to Early Cambrian passive margin, shelf and off-shelf sedimentation. Provenance from a late Mesoproterozoic to Neoproterozoic source

Fig. 2. Interpreted sequence of pre-Carboniferous events in the Sierras de C6rdoba. ultrabasic rocks. Continental sedimentary rocks of Carboniferous, Cretaceous and Tertiary age partially cover this basement (Gordillo & Lencinas 1979). Rift-related Cretaceous mafic rocks (Kay & Ramos, 1996) and Tertiary volcanic centres associated with Andean (Miocene) 'flat-slab' subduction are also emplaced in the basement (Kay & Gordillo 1994). Metamorphic rocks of middle-to-high amphibolite facies are the main and most widespread component of the basement; low-grade metapelites appear only along the western side of the SC, whereas granulite facies meta-basic rocks crop out discontinuously in the central and southeastern areas. Medium-grade para- and ortho-gneisses and schists constitute the dominant lithology (Fig. 1); large massifs composed of garnet-cordierite pelitic migmatites are also characteristic. Subordinate amounts of marble, amphibolite and discontinuous strings of ultrabasic rocks complete the association. Ultrabasic rocks located on the western side of the SC are dominated by harzburgites, chromitites and serpentinites, showing a N-MORB type geochemical signature. These have been interpreted as a disrupted ophiolitic sequence, whereas ultrabasic bodies on the eastern side are composed of lherzolite and websterite with back-arc geochemical characteristics (Escayola et al. 1996).

Igneous rocks of granitoid type occur both as an early, compositionally heterogeneous group of medium-to-small plutons, and as a postorogenic group of large batholiths and plutons (Fig. 1). The former group is associated with the major Early Palaeozoic orogenic events and its characteristics and age are central to the conclusions of this paper. Within it, three subgroups can be recognized. (a) Gla: metaluminous hornblende-biotite orthogneisses, a typical sequence of which crops out in the northeastern sector of the SC. (b) Glb: peraluminous, cordierite-sillimanite monzogranites, associated with the high-grade migmatitic rocks. This group includes the E1 Pil6n complex and the Juan XXIII granite in the NW sector. (c) G2 trondhjemite-tonalite group, commonly developed on the eastern side of the Sierras de C6rdoba (Sierra Chica block), including the Gfiiraldes, La Fronda, San Agustin and Calmayo bodies (Fig. 1) (Caffe & Baldo 1994; Martino et al. 1995; P6rez et al. 1996). Granodiorite plutons emplaced along the western side of the Sierras de C6rdoba, e.g. the La Playa granodiorite (Fig. 1) (G6mez & Lira in press), are also included in this group, since they were emplaced in the same period of magmatic activity and share some geochemical links with the trondhjemite-tonalites (see below).

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Granitoids of the post-orogenic group (G3) are the most voluminous. They were emplaced well after Cambrian and Early Ordovician deformation and metamorphism, in some cases developing contact aureoles in the medium-tohigh grade metamorphic envelope. Examples of these bodies are the Achala and Cerro Aspero batholiths (Fig. 1), a common feature of which is an abundant porphyritic phase with K-feldspar megacrysts (Lira & Kirschbaum, 1990; Rapela et al. 1990). A Rb-Sr isochron age of 358 + 9 Ma (initial SVSr/86Sr=0.7126; M S W D = I . 3 ) has been obtained in the northern sector of the Achala batholith (Rapela et al. 1991). In composition, biotite + muscovite monzogranites predominate over granodiorites and scarce tonalites. These post-orogenic granites are enriched in incompatible elements such as K, Rb and U and LREE, and in some cases in HFS elements such as Nb and Y (Rapela et al. 1990). Post-orogenic activity lasted until the Early Carboniferous~ where the later members of this group were emplaced to the north of the Achala batholith (Capilla del Monte granite, 337 + 2 Ma, Rb Sr isochron, Pankhurst & Rapela, unpublished data) Deformation

Detailed structural studies have been carried out in the SC (Martino et al. 1995; Baldo et al. 1996a and references therein.). However difficulties in correlating structures over long distances and the absence of accurate geochronological constraints have hindered precise understanding of the regional tectonic history. Regional foliations

Three syn-metamorphic regional foliations have been recognized (S1, $2, $3). The first foliation (S1) is usually preserved in the high-grade migmatitic areas as a relict foliation within refractory remnants, or as an internal foliation (Si) in garnet porphyroblasts, often at a high angle to the external foliation. The age and kinematic significance of this deformation remain unknown, although the possibility of a Proterozoic event can not be ruled out (e.g. Baldo et al. 1996). The most widespread foliation in the migmatitic rocks is currently considered to be $2 (Fig. 1). This overprints S1 and is related to the development of large, kilometre-scale recumbent folds with a westward vergence. It is well defined by the alternation of leucosomes and melanosomes in metatexites and more vaguely by disrupted restitic material in diatexites (nebulites). Garnet porphyroblasts, which are very

common in the migmatites, are often concentrated in the leucosomes, suggesting that this deformation took place close to the peak of metamorphism. $2 structures are conspicuous in the G l a metaluminous granitic gneisses of the eastern SC as a regional foliation defined by the orientation of biotite, hornblende and the widespread mafic enclaves of the suite. This foliation shows a fairly regular N-S to NNW-SSE regional trend and dips of 40-50~ Decimetre-scale intrafolial folds of syn-S2 leucosomes are commonly present in the metatexites, suggesting a non-coaxial component. Foliation $2 was generated during the Pampean orogeny (Early to Mid-Cambrian) as evidenced by geochronological work presented below. A third foliation is related to a younger phase of upright folding. These folds show N-S trends and vary from open to very tight isoclinal; in the latter case $2 is rotated into parallelism with the $3 axial plane foliation. Folding related to $3 is accompanied by ductile shearing (StuartSmith et al. 1996). Crenulation and crenulation cleavage displayed by medium-grade schists could also be assigned to this phase. Since $3 deformation affects the G2 trondhjemitic plutons, for which an Early Ordovician crystallization age has been obtained in this work, it is ascribed to the Ordovician Famatinian Orogeny, which is typically developed along the Central Batholithic Belt of the Sierras Pampeanas (Pankhurst et al. this volume). K - A r ages in the range 435-447 Ma for muscovite from a pegmatite swarm associated with $3 and a c. 440 Ma U P b monazite age on a 500 Ma trondhjemitic pluton affected by $3 are consistent with this interpretation (see below).

Shear zones

Eastward-dipping, N-S-trending shear zones are abundant in the SC. Repeated shear deformation, sometimes along the same shear zone, seems to have been very common, and is recorded by the superposition of ductile-to-brittle structures and textures (e.g., foliations, mylonites, cataclasites and breccias). Several of these shear zones are of regional extent and divide the SC into several major longitudinal N-S to N N W SSE narrow domains (Kraemer et al. 1995). The Achala batholith cuts one of these major shear zones. Their importance is further emphasised by the alignment of basic and ultrabasic bodies found along the central part of the SC, considered to be part of a dismembered ophiolitic complex by Escayola et al. (1996). Two main stages of shearing may be postulated:

THE PAMPEAN OROGENY An older stage of syn-metamorphic westward thrusting, close in time to $2 has been recognized (e.g., Martino et al. 1995 among others). It is recorded by a strong E-W stretching lineation in garnet-cordierite gneisses within the shear zones and in mylonites. A subsequent stage of important extensional movements under amphibolite to greenschist-facies conditions has been advocated by Stuart-Smith et al. (1996), probably related to the emplacement of minor, late, peraluminous granitic bodies, such as the Loma Ancha granite in the Sierra Chica (Rio del Suquia section, Fig. 3a). A younger stage of deformation along very steep shear zones is represented by mylonites and local phyllonites displaying relatively flatlying stretching lineations and a dominant dextral sense of movement. This stage is tentatively correlated here with the $3 upright folding referred to above. Older shear zones were reactivated as new vertical wrench shear zones at this stage. Contemporaneous regional metamorphic conditions are not precisely known. Two igneous events separate the stages of shearing referred to above: injection of a swarm of tonalitic dykes and small stocks of the same composition (G2 trondhjemite-tonalite group), well represented in the Sierra Chica area, and intrusion of a younger swarm of large muscovitegarnet-bearing pegmatites, well displayed in the northeastern SC. Contacts of the trondhjemitictonalitic plutons and dykes are usually steep but discordant to vertical $2 structures. They show a weak internal foliation. Pegmatite dykes however are mostly flat-lying and do not show significant internal deformation, except in the La CumbreAscochinga area, where upright folding and longitudinal wrench faulting are strong. Here pegmatites are folded and transposed along the dextral shear zones. These pegmatite bodies have been dated at 447 • 7,442 4- 7 and 435 4- 7 (K-Ar on muscovite, Carap~ shear zone, La CumbreAscochinga area, Fig. 3c, Table l), suggesting an Ordovician minimum age for D3. Consequently, the phase of upright folding and dextral wrench shear could be attributed to Famatinian orogenesis (Ordovician), during which a mobile fold belt developed west of the SC (Pankhurst et al. this volume; Saavedra et al. this volume). Ductile shear zones were probably also active in more recent times. K - A r ages of 365 + 10 and 373 + 11 Ma have been obtained on recrystallized amphiboles from a discordant shear zone in the area of Los Tt~neles (Table 1). These ages are close to the emplacement age of the huge G3 Achaia batholith nearby (Fig. 1). This Devonian stage of deformation is still very poorly understood in detail.

185

P r e v i o u s isotopic data

Reconnaissance geochronological studies of the basement of SC were carried out mostly during the period 1960-1975. The results were compiled by Linares & Gonz/~lez (1990) and include analyses performed by different methods on 334 samples of plutonic and metamorphic rocks; they show an age range from 227 to 1516 Ma. A restricted subset of 83 samples that gave ages older than 570 Ma was used as evidence of a Proterozoic age for the high-grade metamorphic rocks and amphibolite suites, a conclusion that was adopted in most subsequent geotectonic models for this sector of the Sierras Pampeanas (e.g., Gordillo & Lencinas 1979; Kraemer et al. 1995 and references therein). However, close inspection of the 'Proterozoic' subset shows that a large number of these critical data were obtained from (a) Rb-Sr reference wholerock errorchrons for gneisses and migmatites from widely spaced outcrops, assuming initial 87Sr/86Sr ratios (Cingolani & Varela 1975), (b) whole-rock data for single samples, also assuming an initial 87Sr/86Sr ratio, or (c) Rb-Sr whole-rock errorchrons on amphibolites. Typical errors of +20-100Ma were generally reported at a lo- level without adequate goodness-of-fit tests; there is therefore considerable uncertainty in the geological significance of these ages, and a careful re-evaluation of the problem is now needed.

Metamorphic evolution The P - T evolution of the SC is considered here mainly on the basis of information obtained along the valley of the Rio Suquia, in the Sierra Chica (Fig. 3a). Preliminary estimates of the metamorphic conditions of the E1 Pil6n granitemigmatite complex and the low-grade metamorphic rocks from Los T~meles (Fig. 1), are included as well, along with data from other authors for comparison. Thermo-barometric estimates were calculated with the programme "rWEEQU, which is based on an internally consistent thermodynamic data base (Berman 1991). Rio del S u q u i a area

The Rio del Suquia flows through the Sierra Chica, almost normal to the lithological boundaries (Fig. 3a). Very good exposures of metaigneous and metasedimentary rocks are found along this section, and have long been the target of detailed study (Gordillo & Lencinas 1979; Baldo et al. 1996a).

186

C. W. RAPELA E T AL.

Fig. 3. Geological maps of selected areas of Sierras de Cdrdoba. Localities are indicated for the most important samples for which age determinations or microprobe studies have been carried out in the present study. Two lithological domains, referred to as the La Calera and E1 Diquecito groups, can be distinguished, respectively west and east of the major La Estanzuela fault (Fig. 3a). The La Calera Group is composed of high-grade garnet (Grt)-cordierite (Cdr)-sillimanite (Sill) migmatites (metatexites and diatexites) and gneisses, diopside marbles and amphibolites. (N.B: Abbreviations for mineral names throughout this paper are those of Kretz 1983). The El Diquecito Group consists of two lithological

assemblages separated by the E1 Diquecito fault. To the west there is a monotonous sequence of tonalitic orthogneisses, interlayered with minor high-grade pelitic gneisses (Sill + Grt + Kfs), scattered marble and amphibolite lenses. East of the E1 Diquecito fault, the lithology is dominated by a complex sequence of metaigneous rocks, consisting of variably amphibolitized garnet-bearing pyroxene metabasites (Opx +Cpx), tonalitic gneisses, and metasedimentary rocks (Grt + Sill gneiss and migmatite).

THE PAMPEAN OROGENY

187

Table 1. K Ar analytical data for mineral separates Sample

Mineral

40Arrad. (#1 g-1 )

%K

%4~

Age (Ma)

Los Tfineles shear-zone amphibolites TLT- 166 Amphibole TLT- 165 Amphibole

0.68 0.43

10.5993 6.8363

13.18 9.85

365 • l0 373 • 11

Los Tgmeles low-grade metasediments TLT-161 Illite (<0.5 #)

3.82

88.7572

22.27

517 9 19

La Cumbre-Ascochinga and Cuniputo pegmatites 167.4467 LCA- 147 Muscovite 8.77 167.2952 CUN-114 Muscovite 8.50 165.4492 CUN-111 Muscovite 8.50

13.14 12.60 14.95

435 • 7 447 i 7 442 • 7

A penetrative N-S-trending regional synmetamorphic foliation ($2), is common to the rocks of both groups. This foliation is transposed along shear zones, including the La Estanzuela fault zone, where complex multistage ductile to brittle shear movements are recorded. Near the western boundary of the Sierra Chica, late brittle-ductile shear zones and faults with pervasive retrogression ( M s + Chl) become increasingly common. Metamorphic characteristics in the Rio del Suquia area have been determined on a diatexite from the La Calera Group (RSU-65), a metabasite from the E1 Diquecito Group (RSU-9), and a gneiss from the western domain of the E1 Diquecito fault (RSU-21).

Amphibolites are an essential companion lithology to the metasedimentary sequence, not only in the Rio del Suquia area but throughout the SC basement. They have a typical basaltic composition, and their immobile trace element chemistry closely resembles that of N - M O R B (Fig. 4, Table 2). La Calera Group. The diatexites have proved to be very useful for unravelling the metamorphic history of this part of Sierra Chica (Baldo et al. 1996a). They are mostly garnet-rich weakly foliated nebulites, with remnants of wellfoliated ($2) palaeosome and refractory enclaves of quartz, metapsammite, amphibolite, and hypersthene metabasite. Garnet is found as

1~176 I` ', m'pl ibolile'sof the Sierr'as'de (3O'rdoba' ! F

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+8.5 I +5.6 I +7.9 I--

1

0.1

CsRbBaTh U Nb K LaCePb

Sr P NdZrSmEu Ti Dy Y YbLu

Fig. 4. Spider-diagram normalized to N-MORB (Sun & McDonough 1989) for the amphibolites interlayered with the metapelites of the Sierras de C6rdoba. Abundances of the immobile elements abundances, including REE, are very close to N-MORB composition. The depleted mantle signature indicated by the positive cNdt values is also a strong indication that the amphibolites represent original ocean crust basalts.

d~

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THE P A M P E A N O R O G E N Y

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190

C. W. RAPELA E T A L .

megacrysts (c. 2cm) with cordierite coronas, and is particularly concentrated in the leucosome fraction of the palaeosomes, or scattered in the nebulite. A foliation older than $2 is preserved in relict microlithons in the palaeosomes and in psammitic enclaves. These rocks are interpreted as metapelites, as suggested by their high initial 87Sr/86Sr ratios and chemistry (A1203=17.7%; FeO/MgO=2.05; Na20 + K20 = 4.9%) (Tables 2 & 3). Three metamorphic stages have been established on the basis of textural evidence and internal zoning of garnet (Baldo et al. 1996a) (Fig. 5). The first stage (M1), is represented by garnet cores with abundant inclusions of ilmenite, plagioclase, biotite and sometimes rutile. Alignment of the inclusions defines an internal foliation (S1) discordant to that outside the porphyroblast ($2). The second stage (M2) is represented both by the outer parts of the garnets and by the formation of cordierite coronas. The former consist of alternating inclusion-free zones and zones rich in concentrically arranged fibrolite needles (with some biotite). This evidence suggests epitaxial growth of fibrolite during continued growth of the garnet in a migmatitic melt. The border is irregular due to resorption of the garnet by cordierite. The cordieritic coronas consist of a granoblastic assemblage of cordierite, plagioclase, biotite, some Zn-rich spinel ( Z n O = 11.6%), and ilmenite. M2 is thus interpreted as a migmatitic stage, beginning with garnet growth and followed by corona formation. The matrix of the gneisses at this stage consisted for

the most part of quartz, biotite, K-feldspar, plagioclase and sillimanite; relict kyanite has also been found. Finally, a third stage of metamorphism (M3) is represented by local retrogression. Garnet cores are replaced in small domains by a Zn-rich hercynite ( Z n O = 9 % ) , biotite, sillimanite, K-feldspar, magnetite and ilmenite. Cordierite and spinel (+ilmenite) in the coronas, transform into biotite and prismatic sillimanite, in an open system reaction that probably involved K-feldspar from the matrix. The garnets exhibit chemical zonation. Three zones can be distinguished (Fig. 5, Table 4). The inner one, almost coincident with the inner textural zone, does not show any clear zonal pattern. The second, intermediate zone, which is the larger and corresponds to most of the outer textural zone, displays a well developed diffusional zonation of Mg and Fe contents. Ca shows a continuous decrease outwards, with a concavetowards-the-border profile. The third zone is close to the border and is very thin; Ca shows a small increase in this zone. Mn shows no appreciable zonation at all. The diffusional zoning is explained as resulting from resorption of the garnet by cordierite, with FeMg_l exchange at the contact between the two minerals (Tracy, 1982; Spear 1993), during the M2 stage. This zoning was preceded by internal homogenization of the garnet. The Ca profile in the intermediate and the outermost chemical zones suggests the net transfer reaction: 3 anorthite = grossular + 2 kyanite + quartz. Ca decrease in the intermediate zone suggests an earlier predominance of uplift over cooling, whereas the

Fig. 5. Microprobe analysis and textural sketch of a zoned garnet from sample RSU-65 (migmatite, Rio del Suquia area).

THE PAMPEAN OROGENY

~

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192

C. W. RAPELA ET AL.

!

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Fig. 6. TiO2 v. Fe/Mg in biotites from high-grade metamorphic rocks of the Rio del Suquia area, with inferred relationship to metamorphism. Ca increase close to the border suggests that this relationship was subsequently inverted, cooling becoming dominant relative to uplift. A similar thermal evolution is inferred from the composition of the different textural types of biotite (Fig. 6). P - T values for the cordierite corona formation (M2) have been determined in samples RSU-65 (P = 5.7 • 0.4 kbar and T = 820 + 25~ and RSU-21 ( P = 6 . 3 + l k b a r and T = 8 2 0 + 60~ (Fig. 7; Tables 4 & 5). These are probably close to peak metamorphic values, and indicate granulite-facies conditions. The M3 stage took place at P = 4 + 0 . 5 k b a r and T = 7 1 5 + 1 5 ~ i.e. amphibolite-facies conditions. P - T conditions during the earlier stages of garnet growth under subsolidus conditions (M1) could not be determined because of garnet homogenization during M2. The mineral assemblage of the metabasites consists of P1 + Opx + Bt + Qtz + Mag-Ilm + Cpx + Grt + Hb (Gordillo & Lencinas 1979). Peak metamorphic conditions have been obtained for a metagabbro (RSU-9) formed of rather homogeneous plagioclase (An37~4]), orthopyroxene (En40-44; Fe-hypersthene), biotite (XMg = 0.44-0.54), garnet, quartz and amphibole (tschermakitic Fe-hornblende), with accessory magnetiteilmenite, sphene and zircon. Again, textural evidence suggests three stages evolution. The El Diquecito Metabasites.

first is represented by large relict subhedral crystals of plagioclase that show flexure and mechanical twinning. A few large orthopyroxene crystals with undulose extinction and bent biotite flakes (Bt- 1), can be ascribed to this stage as well, which could represent an igneous event. The second stage is a true metamorphic recrystallization. A fine-grained polygonal granoblastic assemblage formed along the boundaries of the relict strained crystals and consists of unstrained plagioclase, orthopyroxene and oriented biotite flakes (Bt-2; XMg=0.44-0.51, TiO2= 3.24.1%, Fig. 6), with some quartz. A weak foliation developed at this stage can be correlated with the external $2. Garnet porphyroblasts also grew during this event, as evidenced by aligned inclusions of biotite, which are parallel to the biotite in the matrix. We interpret this granulitic stage as equivalent to the M2 stage of the neighbouring migmatitic gneisses. Finally, a third stage (probably equivalent to M3) is evidenced by the formation of a late very finegrained assemblage of amphibole, biotite (Bt-3; XMg = 0.520.5-0.54, TiO2 = 3.9-4.0%, Fig. 6), quartz, titanite and ilmenite, along discontinuous planes that transect the former mineral assemblages. Conditions of P = 8 . 6 + 0 . 8 k b a r and T = 810+50~ (Fig. 7) were obtained for the M2 stage, based on a grossular-rich garnet core (retrograde FeMg_l exchange between garnet and matrix biotite took place in the outer part of

THE PAMPEAN OROGENY

193

Leucosome in the stromatites is composed of Qtz + Cdr + Kfs + P1 (An21). The coarse-grained lepidoblastic melanosome contains B t + S i l l (+tourmaline+apatite), with fibrolite along the leucosome-melanosome interface. A polygonal granoblastic aggregate of P1 + Kfs + Bt + Sill + Ilm 4- Crd is typical of the mesosome. Diatexites with nebulitic and agmatitic fabrics have nodular cordierite (5-10 cm), and in some cases grade into porphyritic Sill+ Crd granite. To the NE of the E1 Pil6n granite (see below), the local tonalitic migmatites contain muscovite in their assemblage: C r d + B t + M s + S i l l + Qtz + P1 (An27) -4-Kfs. P-T conditions estimated for the high-grade migmatites are: P = 3 . 9 + 0.6 kbar; T = 684 -4- 60~ (Fig. 7).

Los Thneles low-grade metasediments

Fig. 7. Estimated P-T conditions for the Rio del Suquia area (RSU-65, migmatite; RSU-21, gneiss; RSU-9, metagabbro) and the migmatites from the E1 Pil6n Complex. M2 and M3 conditions were obtained with TWEEQU (Berman 1991). Reactions 1, 1' (Bt + Sil + Qtz = Grt + Crd + Kfs + melt) and 2 were calculated using GEO-CALC (Berman et al. 1987); Crd, anhydrous cordierite; H-Crd, hydrous cordierite. Reactions 3, 3~ and 3" (vapour absent, Bt + Sil + Qtz = Grt + Kfs + melt) were taken from Spear (1993, fig. 9.9). KZ, kinzigites from the Santa Rosa Massif (south of the Achala batholith, Fig. 1); CP, Cerro Pelado, SW of the Achala batholith; ST, migmatite from the Santa Rosa Massif; SC, San Carlos Massif, W of the Achala batholith; YB, Yerba Buena, north of the Achala batholith. the crystals), and matrix granoblastic plagioclase, orthopyroxene and biotite plus quartz (Tables 4 & 5).

Migmatites from the El Pil6n Complex The migmatitic complex of E1 Pil6n in the northern part of the SC (Fig. 3b) is a lowpressure basement block, separated by major faults from other medium-pressure, garnetbearing high-grade rocks. Both the high-grade metamorphic rocks and the granites are affected by ductile shear zones (D3 deformation?). Migmatites are the dominant lithology; stromatite and diatexite are the common types.

Restricted outcrops of low-grade phyllites bounded by N-S Neogene faults occur along the western side of the SC (Fig. 1) (termed the Pocho Phyllites by Gordillo & Lencinas 1979). In the Los Tfineles area, medium-grade mylonitic gneisses (Bt + Ms 4- Sill + Grt + Qtz + P1) are thrust over the low-grade metasedimentary sequence. In this sector the sequence is a rhythmic alternation of metapsammite (SIO2=6973%, A1203=12-13.6%, N a 2 0 + K 2 0 = 4 . 7 5.2%, MGO=1.24-3.26, F e O t = 3 . 7 5 - 4 . 7 7 % ) and metapelite (SIO2=57-62%, A 1 2 0 3 = 1 6 19%, Na20 + K20 = 5.5-6.2%, MgO = 2.663.77%, FeOt = 4.52-7.14%). A laterally very continuous, planar relict stratification (SO) is overprinted by a N10/20E syn-metamorphic S 1 cleavage, subparallel to SO. The S 1 cleavage is locally deformed by Neogene faults, developing kink bands. The metasedimentary sequence is intruded by thin rhyoliticdacitic sills (SIO2=70-72%, A l k = 5-6%). The common mineral association of the phyllites is Qtz + PI(Ab) + Ms(2M1) + Chl(IIb) + Kfs (pyrite, hematite, tourmaline, ankerite and zircon), indicating a low-grade greenschist facies. Peak temperatures of 250-280~ have been determined for the Los Tflneles by chlorite thermometry (method of Walshe 1986). Very lowgrade (sub-greenschist) conditions are indicated.

P - T conditions in other high-grade areas o f the Sierras de C6rdoba P-T determinations from other areas of the SC have been included for comparison in Fig. 7. Baldo (1992) obtained mean values of T = 685~

194

C. W. R A P E L A E T AL.

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THE PAMPEAN OROGENY and P = 4.5 kbar for cordierite-sillimanitegarnet gneisses north of the Achaia batholith. Values of T=725~ and P - - 5 . 8 k b a r were determined by Gordillo (1984) for the San Carlos migmatitic massif, west of the Achala batholith, and T = 730~ and P = 6.5 kbar for the Santa Rosa and Cerro Pelado massifs, south of the batholith. Particularly interesting is the case of the so called 'kinzigites' of the Santa Rosa massif. These are granulitic rocks consisting of garnet, cordierite and quartz with minor sillimanite, plagioclase, orthopyroxene, hercynite and ilmenite, and some retrograde anthophyllite and biotite; they have been interpreted as restites (Gordillo 1984; Martino et al. 1995; Otamendi & Rabbia 1996). Peak values of T = 750 4- 50~ and P = 7 4- 1 kbar are recorded in Otamendi & Rabbia (1996).

The Lower Palaeozoic granitoids Metaluminous series o f the eastern Sierras de C6rdoba ( G l a ) Meta-igneous rocks are very abundant along the easternmost edge of the SC. They are not restricted to this sector: equivalent granitoid

suites less affected by regional metamorphism extend farther north along the Sierras del Norte de C6rdoba, and along the eastern edge of the Sierras Pampeanas south of Santiago del Estero (Lira et al. 1996). In many areas these rocks have been mapped as orthogneisses, as they are affected by the $2 penetrative ductile foliation. In other areas penetrative structures are poorly developed and the shapes of the original plutons and contacts between them are easily recognized. The La Cumbre-Ascochinga cross-section in the NE corner of the SC is a well-exposed example of these sequences (Figs 1, 3c). Here, a series of granitoid plutons was emplaced into a mediumgrade polymetamorphic basement composed of diopsidic marbles, amphibolites, biotite-muscovite schists and tonalitic orthogneisses. Four main plutonic units have been distinguished, all with foliation produced by ductile deformation, and with ubiquitous epidote in the mineral assemblages (P6rez et al. 1996). Some are Late Cambrian/Early Ordovician trondhjemites and are described separately, but the most abundant plutonic rocks, e.g. in the Ascochinga unit, are hornblende-biotite tonalites, biotite tonalites, porphyritic granodiorites and monzogranites (Figs 8, 9). The dominant facies

Strongly PeraluminousUnits High-NaUnits

LEGEND

MetaluminousUnits + Ascochingaunit x Riodel Suquiasection

IEI Pil6n complex:

o Porphyriticmainfacies L9 Leucogranitefacies [] JuanXXIIIgranite

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SanAgustfn A LaFronda 0 Guiraldes

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Ab

Or

Fig. 8. Classification of the Cambrian and Ordovician granitoid rocks of Sierras de C6rdoba. (a) Multi-cationic homologue of the modal QAP triangle, after La Roche (1992); (b) CIPW normative Ab-An-Or granitic rock classification after Barker (1979).

196

C. W. RAPELA E T A L . ' ' ' ' 1 ' ' ' ' 1 ' ' ' ' 1 ' ' '

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Fig. 9. Harker variation diagrams for Cambrian and Ordovician units of the Sierras de C6rdoba. Limits in the K20 plot are from Le Maitre et al. (1989). ASI, alumina saturation index (tool A1203/(mol Na20 +tool K20 + tool CaO)); dashed line at ASI, 1.! is the limit between I- and S-type granites in the Lachlan Fold Belt (Chappell & White 1992). Closed fields correspond to tonalite-trondhjemite units. Symbols as in Fig. 8. is tonalite-granodiorite, with S I O 2 = 6 1 - 6 8 % , ASI = 0.95-1.03, Alk = 3.8-6.5%, M g # - - 3643, K / R b = 1 9 0 - 3 1 0 , [La/Yb]N=4-35, slight depletion in M R E E and small or absent negative Eu-anomalies (Table 2). These rocks carry

abundant disc-shaped hornblende-clinopyroxene gabbroid enclaves (47-51% SiO2), orientated parallel to the foliation. Geochemical trends are typical of medium-K calc-alkaline I-type granites, with slightly peraluminous monzogranite

THE PAMPEAN OROGENY

197

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Fig. 10. (a) Y (ppm) Sr/Y diagram of the granitoid units from Sierras de C6rdoba. Archean High-A1 trondhjemite-tonatite dacite (TTD) field after Drummond et al. (1996). (b) Data for the Sierras de C6rdoba granites plotted on a Yb + Ta v. Rb tectonic discrimination diagram (after Pearce et al. 1984). Symbols as in Fig. 8.

(74-77% SiO2) as the most differentiated facies (Fig. 9). LIL and HFS elements are in the range of arc-related granites (Fig. 10). In the Rio del Suquia section, the high-grade metamorphism and coeval deformation that affected the igneous sequence (M2, $2) obliterate most of the primary igneous fabrics. However, the igneous characteristics of the El Diquecito group located to the west of the La Estanzuela fault (Fig. 3a) are still recognized in the field. This suite is composed of dioritic to tonalitic gneisses ( S I O 2 = 5 6 - 6 6 % , A l k = 3 . 4 ~ . 8 % , Table 2), occasionally containing orthopyroxene in the more basic varieties. Dark, deformed igneous enclaves (SiO2 = 52.6%, Alk =3.5%) are common. Major and trace element geochemistry is also within the range of the medium-K calcalkaline suites (Table 2, Figs 9, 10).

Strongly peraluminous granites ( G l b ) Small-to-medium sized bodies of highly peraluminous granitoids (ASI = 1.20-1.38, Fig. 8) are commonly associated with the high-grade metasedimentary rocks in the migmatitic massifs (Fig. 1). The E1 Pil6n complex and the Juan XXIII cordierite-bearing granite to the west of the Achaia batholith are typical examples (Fig. 3b, Table 2). The E1 Pildn complex, situated in the N W sector of the SC (Figs 1, 3b), is an extraordinary example of the association of such granitoids with low-pressure, garnet-absent high-temperature migmatites. The complex covers about 80kin 2 and is composed of (a) a widespread

porphyritic granitoid, (b) a leucogranite facies of 2.7 km 2 located in the NE part of the complex, (c) four ovoid outcrops of cordieritite, each up to 140m in length, enclosed within the leucogranite, and (d) irregular pegmatoids spatially associated with the leucogranite and cordieritite. Geological, geochemical and isotopic evidence indicates a close genetic relationship between the granites, the cordieritites and the country rocks. The main E1 Pil6n granitoid is a porphyritic siilimanite-cordierite-biotite granodiorite, transitional to monzogranite (SIO2=69-71%, ASI =1.26-1.38, Alk = 6.7-7.2%, Mg# = 46-49, K/Rb = 180-220, REEt = 110-150ppm, [La/ Yb]N=7.5-10.5, Eu/Eu*=0.52-0.63) (Figs 9, 10; Table 2). It has sharp contacts against the high-grade metasedimentary envelope. Country rock xenoliths of all sizes, and widespread clusters of biotite+sillimanite, are common. The leucogranite facies is a sillimanite-cordieritebiotite-garnet medium-grained leucomonzogranite (SiO2 = 71-74%, ASI = 1.23-1.34, Alk = 7.4-8.2%, Mg# = 25-45, K/Rb = 185-240, REEt = 45-90 ppm, [La/Yb]N = 4-5.5, Eu/Eu* = 0.42-0.63). The leucogranite clearly intrudes the main porphyritic granodiorite, carrying irregular-shaped 'xenoliths' of the latter. Cordieritebearing pegmatoids, tourmaline pegmatites and high-silica leucogranites form a different granitoid group closely associated with the leucogranitic facies but with a distinct composition: SIO2=67.5-76%, A S I = I . 1 7 - 1 . 2 1 , A l k = 7 . 5 9.2%, M g # = 31-49, K/Rb = 145-325, REEt = 11-30 ppm, [La/Yb]N = 2.5-4.0, Eu/Eu* = 1.051.40. Orbicular and massive cordieritites are closely associated with the leucogranite facies.

198

C. W. RAPELA ET AL.

The orbicular facies consists of 5-40 cm eggshaped orbicules, comprising an outer aggregate of fresh polygonal cordierite ( M g # = 66) with fibrolite between the crystals and minor quartz and plagioclase (An20_26), and a dark nucleus of orientated biotite, sillimanite, and cordierite • feldspar. The space between orbicules is occupied by a 'leucogranite' matrix composed of quartz, perthitic K-feldspar and plagioclase (An16 20), which in some cases increases in proportion, resulting in isolated orbicules 'swimming' in the leucogranite. The massive cordieritite approaches a composition of almost monomineralic (90-95%) cordierite, and shows a gradual transition to darker varieties with up to 40% of modal biotite. The massive cordieritite (SiO2-- 45-59%, ASI = 6-44, Alk = 0.52%, Mg#=61-67, K/Rb=100-140, R E E t = 10-130ppm, [La/Yb]N= 1.5-3, Eu/Eu*=0.150.9) (Table 2) carries xenoliths of the main facies porphyritic granite, and it displays sharp magmatic-style contacts against both the granite and screen-like exposures of schist (Bio-Qtz-P1Kfs). Dykes of the massive cordieritite intrude both the leucogranite and the schists, indicating that the cordieritite was ductile enough to be mobilized together with the granitic melts. Compared with the other granitoid groups, the G l b peraluminous granitoids are enriched in LIL elements such as Cs, Rb and K, and strongly depleted in Ca and Sr for a given silica content (Fig. 9). They display very low Sr/Y ratios, and in tectonic discrimination diagrams they tend to plot in the syn-collision field (Fig. 10). G2 Group." high-Na intrusive plutons Trondhjemite-tonalite units. These are small to medium size plutons with sharp contacts, mostly located along the eastern side of the SC. All post-date the main $2 structure of the country gneisses and migmatites. Three bodies have been studied in this work, Gtiiraldes (Fig. 3c) and La Fronda in the northeastern sector and San Agustin in the southeast (Fig. 1). Almost all samples have an igneous fabric overprinted by the $3 brittle deformation, producing fracture cleavage and kink structures in the micas (Caffe & Baldo 1994; Martino et al. 1995; P6rez et al. 1996). In the Calmayo and San Agustin bodies, the brittle $3 deformation zones are transgressed by the Achala batholith, so that intrusion of the trondhjemites post-dated the main deformational event D2 but predated intrusion of the Devonian-Carboniferous post-orogenic batholiths (Martino et al. 1995). Abundant modal quartz (>27%) and plagioclase (>40%), and scarce K-feldspar (0-4.5%)

and mafic minerals (<10%), are characteristic of all units, which can therefore be classified as leucotonalites. Apatite, zircon, epidote and monazite are ubiquitous. At least some of the epidote in the Gfiiraldes body is considered to be magmatic (P6rez et al. 1996). Good agreement is obtained between the modal and chemical classifications (both normative and multi-cationic QAP). Using the normative CIPW classification, compositions straddle the trondhjemite and tonalite fields (Fig. 8b), and all contain biotite ( 3 - 8 % ) and muscovite (1-12%). All 13 analysed samples contain 69-75% SiO2 and are peraluminous (Table 3, Fig. 9). The trondhjemite-tonalite unit is enriched in Na20, CaO and Sr, strongly depleted in incompatible elements as K20, Cs, Rb, U, and has very low Rb/Sr ratios compared with the other granitoid groups. The bodies of Gfiiraldes and La Fronda typically have A1203 = 14.6-16.6%, N a 2 0 = 3 . 3 - 4 . 3 % , K20=0.65-1.7%, FeO*/ MgO = 1.5-4; Rb/Sr = 0.015-0.12, [La/Yb]N = 20-55, no Eu-anomaly and low contents of HREE. The San Agustin body has a more extreme composition (Fig. 9), with higher A1203 and Na20, very low K20, Rb and HFS elements, Rb/Sr < 0.03, and highly depleted REE (total REE 20 ppm, [La/Yb]N = 2-5). They plot in the Archaean high-Al TTD field of the Sr/Y v. Y plot (Fig. 10a). Granodiorite units. High-Na plutons crop out in the northwestern sector of the SC, an example of which is the body of La Playa (Fig. l). This is composed of a dominant biotite-muscovite granodiorite facies, and a subordinate facies of two-mica monzogranite (Gomez & Lira in press). It shares with the trondhjemite-tonalite unit a restricted high-silica composition, peraluminosity, and high-Na20, low-K20 characteristics (Table 3, Fig. 9). However, the abundance of other major and trace elements is transitional to those of normal calc-alkaline rocks with the same silica range. On the southern side of the outcrop, there is a small exposure of rather weathered red aplogranite, which appears to overlie the granodiorite and may represent a roof facies.

New U - P b and Rb-Sr ages from the Sierras de C6rdoba Analytical techniques Conventional low-blank U-Pb analysis of small fractions of abraded zircons were performed at the NERC Isotope Geosciences Laboratory, UK (Table 6), using

THE P A M P E A N O R O G E N Y

199

< . ,...~

c,i

r-o e,i

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t-,e-i

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5

4444-44-4.4

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44-+4

-H --H -H -H

-14 ~ -H ~

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ddc~d

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dc~

c~ddd

c~ddd

o

"~" m '.o - -

mc'q

~m

~

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--

-H -n -~ -H

-+4-~

-H~

~ -H -H -H

-H -H -H -H

~

O0

O0 0'~ O0

r'q ,z",

O0 O0 ["'-

2

d

O

0

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eq

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Q,).~

oe-,l

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,~ d

d

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d o

d d d &

d d d &

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C. W. R A P E L A ET AL.

200

Table 7. Rb-Sr whole-rock data for the Sierras de C6rdoba Sample

Rock

Rb

Sr

87Rb/86Sr (%or)

87Sr/86Sr

(87Sr/86Sr)o

Pre-metamorphic metaluminous suite (530 Ma) RSU-71 RSU-72 RSU-73 RSU-74 RSU-75 RSU-77 LCA-117 LCA-118 LCA-119 LCA-120 LCA-121 LCA-122 LCA-123 LCA-125 LCA-126 LCA-127 LCA-128 LCA-129 LCA-130 LCA-131 LCA-132 LCA-133 LCA-138 LCA-139

Gn X Gn Le Le OGn Gn Gn LG LG LG LG X X Gn Gn G G G Gn Gn Gn Gn Gn

25.5 37.7 57.3 95.7 47.4 67.9 60.4 94.6 108.5 115.8 110.3 141.4 39.9 31.4 80.9 194.3 115.3 123.8 113.0 80.2 36.9 32.3 89.8 34.7

236.5 211.7 238,4 196,0 181.4 193.2 317.6 272.3 188.6 178.3 289.2 222.5 471,2 448.6 277.1 173.9 77.2 72.2 146.7 321.7 232.0 241.1 284,4 395.0

0,312 (1.1) 0,515 (0.8) 0.696 1,414 0.757 (0.6) 1.017 0.551 1.006 1.667 1,882 1,104 1,841 0.245 0,203 0.846 3.242 4,336 4.984 2,233 0.721 0,461 0,388 0,914 0,254

0.712420 0.713914 0.715851 0.721788 0.718346 0.720341 0.714692 0.715404 0.721351 0.723185 0.716686 0.722329 0.709988 0.710212 0.714770 0.736125 0.744040 0.748601 0.726262 0.713621 0.712622 0.712286 0.717377 0.708080

0.710065 0.710026 0.710597 0.711103 0.712631 0,712656 0.710532 0.707803 0.708758 0.708968 0.708343 0.708420 0.708136 0,708682 0.708381 0.711630 0.711288 0.710952 0.709390 0.708171 0.709141 0.709354 0.710473 0.706161

96.7 88.7 105.0 113.3 60.6 68,0 68.8 62.1 75.7 89.5 64.1

4.266 3.983 3.152 1.743 8,991 9.715 7,254 6.506 7,202 3,788 7,382

0.746523 0.744665 0.737986 0.726775 0.778867 0.784483 0.770421 0.762861 0.768836 0.744900 0.768852

0.714295 0.714576 0.714175 0.713608 0.710944 0.711089 0.715620 0.713709 0.714430 0.716284 0.713089

133.4 123.4 458.7 173.5 178.7 205.1 45.1 162.0 129.2 178.2 151.6 158.2 78.2 90.4 119.0 148.7 89.9 61.0

2.739 2,781 0.082 1.274 1,956 1.831 0,008 2.549 3,621 2.214 2,062 2,122 0.173 0.083 0.186 2.831 3,431 10.941

0.733727 0.736235 0.709312 0.723773 0.729225 0.727734 0.710322 0.734246 0.741053 0.733227 0.729254 0.730399 0.708194 0.707561 0.708872 0.736932 0.743212 0.794326

0.713037 0.715229 0.708693 0.714150 0.714452 0.713900 0.710263 0.714989 0.713700 0.716502 0.713679 0.714371 0.706887 0.706934 0.707469 0.715548 0.717289 0.711675

Los Thneles phyllites (530 Ma) TLT-158 TLT-161 TLT-162 TLT-163 TLT-293 TLT-294 TLT-295 TLT-296 TLT-297 TLT-298 TLT-299

142.0 121.6 114.1 68.1 185.3 224.5 169.8 137.6 185.4 115.6 160.9

High-grade metamorphic rocks (530 Ma) RSU-63 RSU-65 RSU-69 LFR-85 LFR-88 LFR-89 LCA- 137 MSC-156 MSC-157 MSR-179 MSR-329 MSR-330 MSR-176 MSR-177 MSR-178 PIL-60 PIL-310 PIL-301

PGn PGn Ma Sch Gn Gn Ma Gn Gn Gn Gn Gn Kin Kin Kin Sch Sch X

126.0 118.3 13,0 76.2 119.3 129.6 0.1 142.4 161.2 136.0 107.8 115.7 4.7 2.6 7.6 145.1 105.2 226.5

(1.6)

(50)

(4.5) (9.0) (2.8)

THE PAMPEAN OROGENY Table 7.

201

(continued)

Sample

Rock

Rb

Sr

8VRb/s6Sr (%or)

87Sr/86Sr

(87Sr/86Sr)o

Amph~olites (530Ma) TLT-165 MSC-169 LCA-134 SCS-181

5.4 8,3 6.4 0.3

114.0 225.5 93.0 96.2

0.137 (3.8) 0,106 (2.5) 0.198 0,009 (40)

0.705702 0.707395 0.705985 0.704069

0.704666 0.706591 0.704493 0,704001

84.9 77.2 88.0 82.3 92.9 76.8 65.3 47.4 40.6 43.1 63.6 58.5 51.1 52.0 7.0 3.9 21.2 10.9 8.2 60.1 7.9 12.0 35.9 18.5 22.0 30.4 44.3

6.719 6.353 6.193 5.812 5.300 6.792 8.394 15.241 15,527 (0.7) 16.056 (0.6) 11,620 9,011 10.078 11.343 (0.5) 64,276 (3.0) 23.440 (5.1) 10.409 (1.2) 65.450 (1.7) 17.223 (2,7) 5.174 54.469 (2.7) 60,719 (1.7) 7.917 (0.8) 18,969 (1.3) 9.211 (1.1) 6,486 (0.9) 12,741 (0.6)

0.762553 0.759594 0.758663 0.756937 0.752444 0.762705 0.776317 0.825976 0.828449 0.832665 0.799114 0.781359 0.789282 0.798387 1.182036 0.900929 0.792836 1.195990 0.848303 0.751819 1.138911 1.182347 0.776200 0.861655 0.783773 0.762376 0.809048

0.712567 0.712328 0.712584 0.713694 0.713014 0.712170 0.713870 0.712584 0.712928 0.713213 0.712661 0.714319 0.714304 0.713997 0.703827 0.726538 0.715394 0.709046 0.720163 0.713322 0.733665 0.730598 0.717294 0.720528 0.715244 0.714117 0.714258

0.746925 0.744719

0.712750 0.713800

0.707337 0.707356 0.707427 0.706868 0.713026 0. 763597 0.709362 0.709126 0.708979 0.709076 0.707978 0.709199 0.718227 0,718732 0.718040 0.720017 0.727808 0.790706 1.158180 0.714387

0.706267 0.706308 0.706350 0.706260 0.710652 0. 707853 0.707052 0.707037 0.707099 0.705985 0.706786 0.707137 0.705055 0.705327 0.706119 0.704916 0.706620 0.705600 0.692652 0,706039

El PiNnigneous comp&x (522Ma) PIL-314 PIL-43 PIL-52 PIL-55 PIL-56 PIL-61 PIL-313 PIL-46 PIL-47 PIL-48 PIL-50 PIL-51 PIL-53 PIL-54 PIL-58 PIL-59 PIL-302 PIL-303 PIL-309 PIL-311 PIL-312 PIL-007 PIL-57 PIL-304 PIL-305 PIL-376 PIL-377

PG PG PG PG PG PG LG LG LG LG LG LG LG LG MC MC MC MC MC MC MC BR OC OC OC Peg Peg

194.2 168.7 187.5 164.5 169,4 179.3 186.3 246.9 215.1 236.5 253.1 180.9 176.6 202.2 148.4 30.9 74.7 233.8 47.8 106.0 141.9 241.2 96.6 118.0 68.8 67.7 193.5

Juan XXIIIgranite (?522Ma) MSC-167 MSC-168

Gn Gn

159.2 145.3

100.6 101.5

4.594 4,156

436.2 447.8 442.0 428.7 332.2 53.4 357.1 401.5 389.2 384.6 453.3 406.0 127.5 125.5 133.5 120.0 103.6 20.5 7.5 148.9

0.160 0.157 0.161 0.091 0,355 8. 325 0.345 0,312 0,281 0,462 0.178 0.308 1,967 2.002 1.780 2.255 3.164 12.709 69,520 1.247

Post-metamorphic high-Na suite (470 Ma) LFR-81 LFR-82 LFR-83 LFR-84 LFR-90 LFR-92 LCA-141 LCA-142 LCA-143 LCA-144 LCA-145 LCA-146 LPL-149 LPL-150 LPL-151 LPL-152 LPL-153 LPL-154 LPL-155 RSU-64

Tr Tr Tr Tr Tr Ap Tr Tr Tr X Tr Tr Gd Gd Gd Gd Gd G G To

24.1 24.2 24.6 13.4 40.7 152.8 42,6 43.3 37.8 61.4 27.9 43.2 86.6 86.8 82.1 93.4 113.1 89.5 173.1 64.1

(1.1) (1.1) (1.1) (1.6) (0.7)

(1.2) (2.8)

202

C. W. RAPELA E T AL.

Table 7. (continued)

Sample

Rock

Rb

Sr

Post-metamorphic high-Na suite (470 Ma) (continued) RS U-68 To 82.1 147.3 RSU-70 To 101.2 129.9 RSU-79 To 56.0 215.2 RSU-80 To 56.2 211.1 MSR-180 Ga 57.9 621.4 SCS- 182 Ap 5.7 480.2 SCS-183 Ap 18.6 410.5 SCS- 184 Tr 23.0 805.2 SCS- 185 Tr 18.9 747.7 SCS- 186 Tr 18.0 760.1

STRb/86Sr (%~) 1.614 2.255 0.752 0.771 0.270 0.034 0.131 0.082 0.073 0.068

(3.4) (1.3) (1.1) (1.3) (1.3)

87Sr/86Sr

(87Sr/86Sr)0

0.716358 0.720239 0.710068 0.709968 0.705877 0.705539 0.705080 0.704333 0.704417 0.704499

0.705552 0.705136 0.705030 0.704808 0.704071 0.705308 0.704202 0.703781 0.703927 0.704042

Ap, aplite; BR, biotitic restite; G, granite; Ga, gabbro; Gd, granodiorite; Gn, gneiss; Kin, kinzigite; Le, leucosome; LG, leucogranite; Ma, marble; MC, massive cordieritite; OC, orbicular cordieritite; OGn, orthogneiss; Peg, pegmatite; PG, porphyritic granite; PGn, paragneiss; Sch, schist; To, tonalite; Tr, trondhjemite; X, enclave. Rb/Sr by XRF at British Geological Survey (M. N. Ingham), 2or = 1.0% except where stated; Rb and Sr in ppm; 2c~ error on 87Sr/86Sr mostly <0.02%.

procedures as in Vaughan & Millar (1996). To ensure dating of the main crystallization, as far as possible, only high-quality acicular zircons were picked, or more equant grains if they were perfectly euhedral. Rb-Sr and Sm-Nd whole-rock analyses were carried out in the same laboratory (Tables 7 & 8), using standard techniques most recently described in Pankhurst & Rapela (1994). Errors quoted on determined ages are 2~ throughout. Some samples were also analysed for U-Pb using the SHRIMP ion microprobe at the Research School of Earth Sciences, Canberra (Tables 9 & 10). Zircons were hand picked from mineral concentrates, mounted in epoxy together with chips of reference zircons, and polished. Cathodoluminescence images obtained prior to anlysis were used to target the magmatic rims of euhedral grains. U-Pb analyses were carried out during using SHRIMP II, each spot analysis consisting of six scans through the mass range. The data were reduced in a manner similar to that described by Compston et al. (1992) and Williams & Claesson (1987), with augmented uncertainties for the isotope ratios using software written by T. R. Ireland (see Muir et al. 1996). The Pb/U ratios were normalised relative to a value of 0.1859 for the 2~ ratio of the AS3 reference zircon, equivalent to an age of 1099.1 Ma (Paces & Miller 1993). Monazite grains were also hand picked from two concentrates and mounted in epoxy together with a reference monazite from the Maffra Adamellite. The SHRIMP monazite dating techniques, using SHRIMP i, followed those given in Ireland & Gibson (1998) and Williams et al. (1996). U/Pb and Th/Pb ratios for the unknowns were normalised relative to 0.0675 and 0.02105, respectively for the Maffra monazite (see Ireland & Gibson 1998). As with the zircon SHRIMP data, uncertainties for individual analyses (ratios and ages) are reported at the lcr level, but errors in calculated weighted mean ages are 95% confidence limits.

Results

R b - S r data proved to be inadequate for directly establishing the ages of either the pre-M2 granitoids or the high-grade m e t a m o r p h i c rocks. The former were dated by conventional U - P b zircon analyses from three localities (Table 6), plotted in Fig. 11. The zircons from sample LCA-131, a gneissic h o r n b l e n d e - b i o t i t e granodiorite from the G l a metaluminous suite of La C u m b r e Ascochinga, gave three concordant analyses, with one more falling at slightly higher 2~ 235U ratio, possibly due to minor inheritance. The crystallization age was estimated in two ways: (i) a fit to the three most concordant fractions, forced through zero, gave an intercept age of 528 + 1 Ma, and (ii) the weighted m e a n 2~ age for all four fractions yielded an indistinguishable result of 529 4- 2 Ma. Two fractions from a garnet-cordierite orthogneiss in the Rio de! Suquia section (RSU-77) and two from the main intrusive phase of the Sierra N o r t e de C d r d o b a (GTL-II) were insufficient to define independent ages, but when regressed together gave a straight-line fit through zero with an intercept of 533 + 2 Ma. A l t h o u g h these ages are barely within analytical error, we consider that they are close e n o u g h to establish the emplacement of the G la metaluminous suite within narrow limits at, say, 530 4- 3 Ma. The high-grade m e t a m o r p h i s m was dated by S H R I M P U - P b analysis of m e t a m o r p h i c m o n azite from the same garnet-rich migmatite sample in which P - T conditions for M2 were determined (sample RSU-65, Fig. 7); 17 out of 19 spots analysed give a weighted mean 2~

THE PAMPEAN OROGENY

203

Table 8. Sm-Nd whole-rock data for Sierras de C6rdoba Sample

Sm

Nd

1 4 7 S m / 1 4 4 N d 143Nd/|44Nd

(143Nd/144Nd)o eNdt

Tdm#

57.05 32.71 44.08 2.40 20.86 19.42

0.1206 0.1355 0.1306 0.0885 0.1091 0.1056

0.512180 0.512131 0.512075 0.512021 0.512103 0.512104

0.511761 0.511660 0.511622 0.511714 0.511724 0.511737

-3.8 -5.8 -6.6 -4.8 -4.5 -4.3

1541 1685 1740 1610 1595 1575

34.35

0.1185

0.512021

0.511609

-6.8

1756 1686 1617 1597 1493 1591 1587 1554 1573 1696 1761 1822 1477

Pre-metamorphic metalummous suite (530Ma) RSU-73 RSU-75 RSU-77 LCA-122 LCA-125 LCA-131

Gn Le OGn LG X Gn

11.38 7.33 9.52 0.35 3.77 3.39

Los Tune~s Phyllites (530Ma) TLT-294

6.74

High-grademetamorphic rocks(530Ma) RSU-65 LFR-89 MSC-157 MSR-179 MSR-329 MSR-176 MSR-177 MSR-178 ACH-38 PIL-301 PIL-60 PIL-310

PGn Gn Gn Gn Gn Kin Kin Kin Sch X Sch Sch

9.40 8.76 4.04 9.79 11.62 3.26 2.04 2.63 8.05 7.06 9.44 6.58

50.33 41.39 19.58 48.00 63.12 13.06 9.13 8.78 41.32 34.75 42.66 31.55

0.1130 0.1279 0.1248 0.1233 0.1113 0.1510 0.1350 0.1808 0.1177 0.1227 0.1338 0.1261

0.512052 0.512153 0.512156 0.512222 0.512113 0.512254 0.512221 0.512367 0.512062 0.512032 0.512026 0.512242

0.511660 0.511709 0.511729 0.511794 0.511727 0.511730 0.511752 0.511739 0.511653 0.511606 0.511561 0.511804

-5.8 -4.8 -4.7 -3.2 -4.5 -4.4 -4.0 -4.3 -5.9 -6.9 -7.7 -3.0

2.83 3.52 2.47

8.39 9.62 6.38

0.2037 0.2213 0.2336

0.512950 0.513128 0.513205

0.512243 0.512360 0.512394

5.6 7.9 8.5

26.08 28.91 27.89 11.45 4.73 2.28 32.10 59.79 3.99

0.1282 0.1279 0.1290 0.1558 0.1652 0.1366 0.1139 0.1191 0.1771

0.512143 0.512137 0.512145 0.512202 0.512251 0.512142 0.512052 0.512061 0.512261

0.511705 0.511700 0.511704 0.511669 0.511686 0.511675 0.511663 0.511654 0.511655

-5.1 -5.2 -5.l -5.8 -5.5 -5.7 -6.0 -6.1 -6.1

1632 1639 1633 1682 1658 1674 1691 1704 1702

37.92

0.1233

0.512095

0.511667

-5.7

1676

0.0911 0.0973 0.1226 0.2622 0.0982 0.1309 0.1468

0.512068 0.512079 0.512489 0.512864 0.512369 0.512485 0.512474

0.511788 0.511741 0.512111 0.512057 0.512067 0.512082 0.512022

-4.8 -4.2 1.5 0.4 0.6 0.9 -0.2

1573 1570 1069 1160 1144 1119 1217

Amph~olites (530Ma) LCA-134 TLT-165 SCS-181

El Pil6n igneous comp~x (522Ma) PIL-43 PIL-52 PIL-56 PIL-51 PIL-62 PIL-59 PIL-302 PIL-007 PIL-377

PG PG PG LG LG MC MC BR Peg

5.53 6.12 5.95 2.95 1.29 0.51 6.05 11.78 1.17

Juan XXIII Granite (?522Ma) MSC-168

G

7.73

Post-metamorph# h~h-Na suite (470Ma) LFR-81 LCA-143 LPL-149 LPL-154 RSU-79 MSR-180 SCS-185

Tr Tr Gd G To Ga Tr

2.46 2.77 4.25 0.88 8.07 5.33 0.45

16.33 17.21 20.96 2.02 49.70 24.61 1.86

Key as for Table 7. Analyses by mass-spectrometric isotope dilution (Sm and Nd in ppm); 2a error on mNd/144Nd < 0.01%. Tdm# is multistage model age (DePaolo et al. 1991).

age o f 522 4- 8 M a (Fig. 12). T h e 2~ ( n o t plotted) s h o w m o r e scatter, b u t are tially c o n c o r d a n t ; 15 o f the p o i n t s give 7 M a , w h i c h overlaps the 2~ age error. T h e m o n a z i t e is c o n s i d e r e d to have

ages essen534+ within grown

t o g e t h e r with the cordierite c o r o n a s at the t h e r m a l p e a k ( T = 820-I-25~ at the onset o f the M 2 m i g m a t i t i c stage in pelitic protoliths. These ages, especially t a k i n g into a c c o u n t the results p r e s e n t e d below for the G l b g r a n i t o i d

204

C. W. R A P E L A E T AL.

Table 9. Summary of ion microprobe U-Pb results for zircons Spots

U (ppm)

Th (ppm)

Th/U

Pb* (ppm)

2~176

RSU-65 1.1 2.1 3.1 4.1 5.1 6.1 7.1 8.1

735 381 1050 561 418 1377 346 348

180 58 130 329 183 84 182 61

LCA-142 1.1 2.1 3.1 4.1 5.1 5.2 6.1 7.1 8.1 9.1 10.1 10.2 11.1 11.2 12.1 13.1 14.1 15.1 16.1 17.1 18.1

34 319 489 415 354 168 381 417 417 312 378 721 203 244 285 326 413 379 345 313 526

22 62 21l 153 76 30 168 234 112 26 58 494 169 180 181 91 156 138 256 91 264

2~

0.25 0.15 0.12 0.59 0.44 0.06 0.53 0.18

77 55 144 137 61 127 97 33

0.000036 0.000059 0.000084 0.000086 0.000093 0.000040 0.000005

0.1067 + 0.0023 0.1505+0.0059 0.1442 4- 0.0079 0.2174 4- 0.0046 0.13954-0.0078 0.09924-0.0037 0.2638 + 0.0056 0.0938 4- 0.0018

654 4- 13 9044-33 869 4- 44 12684- 24 8424-44 6104-22 15094- 29 578 4- 10

0.65 0.19 0.43 0.37 0.21 0.18 0.44 0.56 0.27 0.08 0.15 0.69 0.83 0.74 0.64 0.28 0.38 0.36 0.74 0.29 0.50

5 43 68 58 47 34 54 64 56 37 50 136 31 36 41 44 56 51 53 42 76

0.001209 0.000010 0.000099 0.000092 0.000092 0.000087 0.000048 0.000096 0.000124 0.000076 0.000233 0.000017 0.000081 0.000059 0.000134 0.000160 0.000169 -

0.07164-0.0021 0.0823+0.0013 0.08044-0.0030 0.08074-0.0019 0.0807 + 0.0024 0.1216+0.0023 0.0805 4- 0.0022 0.0848 4- 0.0022 0.0806 4- 0.0020 0.07494-0.0045 0.08164-0.0014 0.1010 4- 0.0081 0.08014- 0.0028 0.0790 4- 0.0025 0.07844-0.0014 0.0797-t-0.0015 0.0791 4-0.0025 0.07884-0.0029 0.0802 • 0.0014 0.07894-0.0023 0.08144-0.0025

446• 13 5104-7 4994- 18 5004- 11 500 4- 14 7404- 13 499 • 13 524 4- 13 500 4- 12 466• 5064-9 620 4- 48 497 4- 17 490 + 15 4874-8 4944-9 4904- 15 4894- 17 497 4- 8 4894- 14 505 4- 15

Ages (Ma) 206pb/238U

207pb/206U

958 • 41 909 4- 1 6 16674- 11 817+36 1498 -4-22

signifies no 2~ detected. Uncertainties given at the hr level. Correction for common Pb made on the basis of extrapolation to concordia along a mixing line with common Pb, following Tera & Wasserburg (1972), as outlined in Compston et al. (1992).

-

suite, confirm that the M2 m e t a m o r p h i c event followed very closely after e m p l a c e m e n t of the m e t a l u m i n o u s granitoid suite. Zircon in this sample (small clear globular grains) is entirely inherited, and the S H R I M P data show clusters of ages at c. 1400Ma, 8 0 0 - 1 0 0 0 M a and 6006 5 0 M a (Fig. 12), revealing a p r e d o m i n a n t l y M e s o / N e o p r o t e r o z o i c provenance. R b - S r data for several widely spaced samples from the regional schist outcrops are also shown in Fig. 13a, the majority being derived from the section at Los Tfineles. As is c o m m o n with such low-grade rocks, the data show considerable scatter. Eliminating the most a n o m a l o u s data point (for TLT-298) reduces the M S W D substantially from 45 to 25, and taking this degree of scatter into account we obtain an imprecise age of 525 ~ 18 Ma. Despite the uncertainty, this

is consistent with the age of the m a j o r event indicated by the G1 granitoids a n d gneisses, and suggests that the high- and low-grade metam o r p h i c events could have been c o n t e m p o r a neous. A similar result was obtained from lowgrade schists in the F a m a t i n i a n belt ( P a n k h u r s t et al. this volume). The G lb p e r a l u m i n o u s rocks of the E1 Pil6n complex have been successfully dated by both U - P b and R b - S r methods. Figure 14 shows a U - P b concordia plot for zircon fractions from a sample of the m a i n porphyritic granodiorite. One is almost c o n c o r d a n t , but two others show discordance typical of inheritance from crustal sources (one m a r k e d l y so). Since the complex is regarded as largely anatectic, this is not surprising. A fourth fraction is discordant and obviously not collinear with the others, reflecting the

THE PAMPEAN OROGENY

e

l

e

205

l

~

~

elm

o~

e l e l e l e l ~ e l ~ e l ~ e l e l ~ e l e l e l e l e l e l e l

e l e l ~ e l ~ e l e l e l e l e l e l ~ e l e l ~

el el

-~

,.-... exl tr

C,,.~ ,-.., ,..-- t"--I ,....~ t"-,I r

~',--:.el

.

.

t"--I r

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tg'3 t'-,,I ~'-'4 ~"q r

t'---I ~ 1

eq

r

el,.-

Q..) ',~' "~" ee~ r

~"

"~" q"~ '~-'t" t " - q'3 "~" q'3 "~" q ~

"~1" '-t'5 "~"

0.66 I

0.68

I

0.70

I

I

'

I

i

LCA-131

(a)

/

'

540

(b)

55o

0.089

0.08; Intercept (3 points: MSWD= 1.5)

-"

5

0.087 Mean 2~ age 530//',~fl/ (4 points) / ~ / 5 2 9 + 2 Ma p /

-

23

oo r r

0.08,c

_

-

o

0.083

~ /." ~'./ / /~" /~

-

51~ /

/

/

r

/ 0.62

530~4~ ' ,~'--

0.085

/'" /

51o/

~, /

//t //

500 /

."

[ /

- / / / ' /

/

GTL-II(shaded)& RSU-77(unshaded) _

f

,

0.083

,

I 0.64

,

0.081

(4p~

I 0.66

I

I 0.68

I

I 0.70

2o7pb/235U

i

0.72

Fig. 11. Concordia plots of U-Pb zircon data for the pre-metamorphic calc-alkaline suite of the Sierras de C6rdoba. (a) Metaluminous gneiss from La Cumbre-Ascochinga section (LCA-131), all four analysed fractions being essentially concordant. (b) Orthogneiss from Rio del Suquia (RSU-77) and hornblende-biotite granodiorite from Sierra Norte de C6rdoba (GTL-II); one fraction of each shows variable recent Pb-loss, but the other two are concordant. The age is derived by forcing regression for all data through the zero age point. Error ellipses are 2-~.

0.080

....

, ....

i ....

i ....

, ....

, ....

, ....

, ....

, ....

, ''

'

|

I

o Zircon 2 further points plot with 2~176 ages of 1 3 5 0 and 1 4 3 0 M a -

0.075

"N,.,, -'N=. ~ _ 1000 " ~ - - - - 0 - - , "o,.. 0070 c~ 900~__.~ o.o6

-

80o

i I ~l ~ I ~

+

9 Monazite n= 17 522+8Ma M S W D = 1.2

"--....

-

-

-

o.ooo

o.o55

RSU-65

0,050

. . . .

5

I

6

I

,

,

I

I

7

I

I

2381 I 2O6p / b I

I

I

8

. . . .

I

9

. . . .

I

10

. . . .

I

11

. . . .

I

12

. . . .

I

13

. . . .

I

14

. . . .

5

Fig. 12. Tera & Wasserburg (1972) concordia diagram for U-Pb SHRIMP data for a cordierite-garnet migmatite from the Rio del Suquia (RSU-65; see also Fig. 3). Analyses are plotted, uncorrected for common Pb, with 1~ error bars. The monazite data (filled squares) form a reasonably tight cluster. Excluding two points which appear to have suffered Pb-loss (marked with crosses), these give a weighted mean 2~ age of 522 :t: 8 Ma (95% confidence limits), interpreted as the age of the main metamorphic event (M2) of the Sierras de C6rdoba. The dashed line is a mixing line to 522 Ma common Pb. Small rounded zircons from the same sample (open symbols) give ages reflecting Neoproterozoic inheritance (four have 2~176 ages of 800-1000 Ma, two give older ages, and one gives c. 600 Ma).

THE PAMPEAN OROGENY

87Sr/86Sr (a) Taninga - Los T~neles - El Pilbn 0.80. Schists and Phyllites i

9 Los Tuneles 9 El Pil6n district zx Enclave in I El Pil0n granite I

0.78. ' 0.76-

excluded point

0.74

,/// /f"

Age 525 + 18 Ma I.R. 0,7141 _+0.0012 MSWD = 25 (Enhanced Errors)

~/C~e~ / # ~ " 9

0.72.

//=r

//-

87Rb/86S r I

0

,

2

87Sr/86Sr

I

,

4

I

~

6

I

,

8

1.10 1.00

112

,

./

9 l./// ,//

I.R. 0.7136 _ 0.0005 MSWD = 2.3 n-18 / /1 - " /~,/ _~

0.90

110

(b) El Pilbn Igneous Complex

9 Granular Granite 9 Massive Cordieritite 9 Orbicular restite 9 Pegmatite

1,20

0.80

/ "=~- ~" f

//'~ _ // ,jr,/-mj"

/~//"//" ///

o D t, /,

~

PorDhvritic Granite - Nodular Cordieritite Schist Enclave in PG Local

Schists 87Rb186Sr

i

I

10

i

I

20

i

I

30

i

I

40

i

I

50

i

I

60

i

I

70

Fig. 13. Rb-Sr isochron diagrams. (a) Low-grade metasedimentary rocks from the Sierras de Cdrdoba. These scatter considerably about an errorchron which gives an imprecise age comparable with that of the M2 anatectic event. (b) The E1 Pil6n complex: the units representing the anatectic equilibrium define an isochron with an age virtually indistinguishable from the U-Pb age of Fig. 13.

probable complexity of this inheritance and/or subsequent Pb-loss. A least-squares fit ignoring this last point yields a lower intercept at 523 • (MSWD =2.0), which is taken as the age of emplacement. Rb-Sr data for the E1 Pil6n complex are included in Table 7 and plotted in an isochron diagram in Fig. 13b. This shows a remarkably good fit to analytical points from the leucogranite phase, the massive cordieritite, a restitic biotite-sillimanite kernel, and two samples of associated pegmatites. These data yield an 18-point isochron of 523 + 4 Ma, with an initial 87Sr/86Sr ratio of 0.7136 + 0.0005 ( M S W D = 2 . 3 ) . The data for the porphyritic granite, a schist enclave within it, and samples of the local schist envelope all fall close to this isochron, though not precisely on it, indicating strong genetic links between all phases of the complex. Although beyond the scope of this paper, a self-consistent geochemical model can be argued for the generation of the leucogranite as an anatectic melt of the local schists in the

207

roof zone of the main granodiorite, with the formation of the cordieritite as a peritectic reaction product. We thus consider that the age of the G lb peraluminous granite group at E1 Pil6n must fall within the interval 525-519Ma (maximum U-Pb age for the porphyritic granite, minimum Rb-Sr age for the anatectic rocks). Geochronological studies of the G2 high-Na suite indicate more than one event. The Gfiiraides trondhjemite has given a SHRIMP U-Pb zircon age of 499 + 6 Ma (Fig. 15a), consistent with post-metamorphic emplacement. This zircon age is supported by conventional U-Pb analysis (Fig. 15b): one fraction is concordant at 500+ 14Ma and the others show evidence of inheritance of different ages (from c. 1000 Ma to >2000Ma). The most precise age obtainable is by linear regression of the three fractions with the lowest 2~ ratios, defining a lower intercept of 496-t-2Ma. However, SHRIMP data for monazite from the same sample show the strong influence of younger resetting; the full data set gave a much younger age of 438 -4-9 Ma, with an MSWD of 2.7 indicating excess scatter (2~ ratios, after discounting two higher ages, gave 438-t-7 Ma). The resetting event, at 440Ma or younger, may have been heating associated with the D3 dextral shearing. The La Fronda trondhjemite and the La Playa granodiorite were both analysed by Rb-Sr methods; the low-Rb/Sr rocks were complemented by a muscovite separate and an aplite from La Fronda, and by two samples of the evolved red granite at La Playa. The data are shown in Fig. 15c. Although there is scatter beyond analytical error, this is possibly due to inhomogeneity within the xenolithic granodiorite; exclusion of three samples of this phase results in an errorchron age of 4 7 4 • (initial 87Sr/86Sr= 0.7063, M S W D = 4 . 3 ) . The muscovite-whole rock pair from La Fronda independently give an age of 468 + 6 Ma.

Isotopic constraints The initial 87Sr/86Sr ratios of the G l a metaluminous suite (calculated at 530 Ma, Table 7) mostly vary from 0.708 to 0.713. This wide range indicates considerable crustal input to the calcalkaline magmas, which are thought to have been subduction-related but clearly emplaced into continental crust. This is also borne out by their initial GNd-values of -3.8 to - 6 . 6 (Table 8). The high-grade basement rocks into which they were emplaced, prior to the M2 event, have a slightly stronger crustal signature, with 87Sr/S6Sr

208

C. W. RAPELA E T A L .

0.106 0.102

?

0.098

..Q 0.094

El Pil6n Porphyritic Granite PIL-52

640 620

Intercepts 523 _+2 Ma & 1216 + 77 Ma (3 points: MSWD= 2.0)

0

2

54~ / /

j-

0.086

/

,s i

~o ~ 0.090

.J

I

I

0.086 -

5,

y

0.084-/~/~y

0.082 0.0781 / , 0.60

0.66 0.64

0.68

0.72

0.76

0.80

0.68 0.84

,

I 0.88

2oTpb/ 23sU Fig. 14. Concordia plot for U-Pb zircon data from the main (porphyritic) granite of the E1 Pil6n complex. Three of the points are collinear, with a lower intercept representing igneous emplacement and a reversely discordant upper projection signifying inheritance from an older crustal protolith. The fourth point appears to have suffered more recent Pb-loss and is ignored.

ratios at 530Ma generally >0.710 and cNd530values o f - 3 . 0 to -7.7. The chief exceptions are the kinzigites of the Santa Rosa massif, which are isotopically more primitive with respect to their 87Sr/86Sr530 ratios of ~0.707, although their cNd530 values (-4.0, -4.3) fall within the range of the other high-grade rocks. The low-grade phyllite sequence of the western margin of the SC has isotopic compositions at 530 Ma similar to the more evolved high-grade rocks, and was of continental provenance. It is feasible that the protoliths of the high- and low-grade rocks were identical. Further evidence for the petrogenesis of the El Pildn complex, and for the composition of the continental basement of this area, is given by the Sr- and Nd-isotope compositions listed in Tables 7 and 8. These show uniformly radiogenic initial 87Sr/86Sr ratios of, most significantly, 0.7125-0.7135; the values that fall outside this range correspond to samples with very high Rb/Sr ratios, whose time-corrected compositions are thus subject to large error magnification. The uniformity is evident from the cNd522 values which fall in a narrow range of -5.5 to -6.1 for the leucogranite and cordier-

itites of the complex, with a slightly lower value of -5.1 for the main porphyritic granodiorite. The Sr and Nd compositions are consistent with the crustal melting model - in fact the anatectic rocks have more uniform cNdt values than the basement rocks on a regional scale, possibly due to the homogenizing effects of anatexis on a large scale. The post-metamorphic G2 granitoid group shows variable composition. The bodies at Gfiiraldes and La Fronda show important crustal influence, consistent with their mica-bearing mineralogy, whereas the San Agustin trondjhemite and the other undeformed Ordovician plutonic rocks of this group are much more primitive (cNdt values around zero and, for the former, initial 87Sr/86Sr ratios close to 0.704). The nature of the deep continental crust of the SC is best exemplified by the Nd model ages of these rocks (Table 8, Fig. 16). These were calculated according to the model of DePaolo et al. (1991) which assumes a major fractionation of Sm/Nd when the crustal reservoir is first formed from depleted mantle, followed by a single stage of crustal residence up to the time of magma generation/metamorphism. Despite the

THE PAMPEAN OROGENY 0.090

-''

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209

9 Zircon n=18 4 9 9 +- 5 Ma LM S W D = 0.7

I . . . .

I . . . .

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(a) 9

Monazite

~

n= 16

,T, ~ T ', '.

4 3 8 +9 Ma M S W D = 2.7

,

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0.065 0.060

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_

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, , I , , , , I , , L , l , , , l l , , , , I

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.

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-

A~te

~.~.

Age 474 + 6 Ma_ MSWD 4.3 Enhanced Errors

0.72

'" Pb/'" U k60

~

I

0.62

,

I

0.64

,

I

0.66

,

I

0.68

,

18

o: La Playa Whit~ Playa Red La Fronda

0.74

"

17 Plutons

-g'90 ,,2'~

0.079

I , , , ,

16

LaPlayaandLaFronda

0.82.

496 + 2 Ma

0.081

(C)

14

87S 86Sr

~._ 5./

L C A -(3 1 4points): 6

. . . .

13

87Rb/86Sr

I

0.70

0.70

0

'

'

:

:

:

5

:

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=

:

=

10

~

=

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=

15

Fig. 15. Geochronological data for Famatinian plutons of the Sierras de Cdrdoba: (a) U-Pb SHRIMP data for the Gfiiraldes trondhjemite, plotted as in Fig. 12. The zircon analyses (open circles, excluding two for inherited cores marked with crosses) define a weighted mean 2~ age of 499 4-6 Ma; the monazites (filled circles) record resetting by a subsequent event at c. 440 Ma or younger. The dashed lines are mixing lines to contemporaneous common Pb. (b) Conventional U Pb zircon analyses for the Gfiiraldes trondjhemite: one concordant point and two fractions showing simple inheritance defined by a chord with the lower intercept of 496 :k 2 Ma; the fourth point, although the most precise analysis, is anomalous in showing a different age of inheritance. (e) Rb-Sr data for the La Fronda trondhjemite and La Playa granodiorite.

simplicity of these assumptions, the model yields consistent late Palaeo-to-Mesoproterozoic crustal ages of 1500-1800 Ma for the majority of the rocks, with the late high-Na magmas giving younger model ages of 1100-1200 Ma. There is a tendency for such model ages to over-estimate the age of major rock-forming orogenies in continental crust; data from the Grenville Province of Canada (over 100 analyses from the literature calculated in the same way; RJP unpublished) show that the granites have model ages peaking at 1300-1600Ma, with orthogneisses c. 100Ma older. We conclude that the major rock-forming event in the crystalline basement of the SC was probably mostly Palaeoproterozoic or Mesoproterozoic in age.

Discussion The Cambrian orogeny of the Sierras de C6rdoba Tectonothermal and magmatic evolution. New precise U - P b and Rb Sr ages for several metamorphic and plutonic units considered to be among the older components of the basement of the SC are of Early Cambrian age (time-scale of Gradstein & Ogg 1996). The analysed rocks include the pre-metamorphic metaluminous orthogneisses of the eastern sector, typical medium-pressure migmatites, the low-pressure high-grade rocks and related strongly peraluminous granites of the E1 Pil6n complex, and

C. W. RAPELA ET AL.

210 [ ] G1 Metaluminous Suite [ ] G2 Peraluminous Suite (El Pil6n Complex) 9 High-Grade Basement [ ] Low-grade Basement [ ] G2 High-Na Suite

n

...... =

|

,

*

|

i

|

n ,lL i

u

,

i

,

i

.

..... i

i

i

I

O

O

O

CD

O

O

O

O

O

CO

CO

O

04 x'-

'~" ,e'-

CO ~

CO x--

~ 04

04 04

i

Tdm* Ma

Fig. 16. Histogram of multi-stage Nd model ages for igneous and metamorphic rocks from the Sierras de Cdrdoba, calculated assuming a previous crustal stage originating in depleted mantle at the time shown. Data from Table 8. (within broad error limits) the low-grade Los Tfineles phyllites of the western border of the SC. These consistent results obtained on different lithologies and by different methods indicate that the main geological evolution of the basement occurred during the early Palaeozoic. As current regional correlations and geotectonic models have assumed a Proterozoic age for the Pampean metamorphism (e.g. Kraemer et al. 1995), the new results may have a profound effect on the interpretation of pre-Andean regional geology and particularly on models for the assembly of the supercontinent of Gondwana. Although individual P - T paths are not available, except for the two P - T determinations (M2 and M3) of the La Calera diatexite, the overall pattern of the metamorphic history of the SC migmatite gneisses is one of a clockwise P T path, with a significant stage of isothermal decompression from granulite-facies conditions ( T > 800~ M2 stage) (Fig. 7). The M2 highgrade event reached temperatures high enough to cause widespread anatexis of the metasedimentary rocks. Mineral paragenesis in the migmatites indicate that temperatures largely surpassed the granite minimum-melting curve, and that incongruent melting of biotite commonly occurred during this process. For the Rio del Suquia section, the inferred path between M2 and M3 corresponds to a decompression-uplift of approximately 13km and cooling through 100~ At the end of this stage, at relatively low pressures and lower temperatures ( T = 700-750~ significant melting of the gneisses led to the local formation of

anatectic granitoids, such as those of the E1 Pil6n complex. Uplift of a gravitationally unstable thickened continental crust by a combination of erosion and tectonic denudation is typical of continental collisional settings (e.g. Thompson & Ridley, 1987; Jamieson 1991). Tectonic denudation is also invoked by Harley (1989) and Anovitz & Chase (1990) among others to explain isothermal decompression of granulites. Crustal thickening drove supracrustal rocks and plutonic rocks of the E1 Diquecito group to depths equivalent to pressures probably in excess of 7-8kbar, i.e. 26 30km. Tectonic denudation probably took place along extensional detachments and the La Estanzuela fault was probably one of these faults, albeit masked by subsequent reactivation.

The Pampean Orogeny o f N W Argentina A Late Neoproterozoic-Cambrian sedimentary and tectono-metamorphic sequence of events known as the Pampean Cycle is well established in accounts of the Eastern Cordillera, Puna and northwestern Sierras Pampeanas (Acefiolaza et al. 1990). Due to the high structural level of basement exposures in this sector, chronological and stratigraphical relations have been established on the basis of geological and palaeontological evidence. The oldest unit of the basement is the Puncoviscana Formation, first described in the Eastern Cordillera (Turner, 1960), which is mainly composed of a thick sequence (>2000 m) of lowgrade metapelites and turbidites with well preserved primary sedimentary structures. This tightly folded series is unconformably covered by shallow marine platform sediments of the Mesdn Group, that carry poorly preserved late Mid- to Late Cambrian trilobites of Laurentian affinities (Bordonaro 1992). Inferred equivalents of the Puncoviscana Formation are a widespread component of the basement of NW Argentina (Puncoviscana Formation s.l., Toselli 1990), and they extend southwards through the Famatina System as far as 29~ forming the host country rocks of the Ordovician granites (Durand & Acefiolaza 1990). Importantly, the metasedimentary rocks of the northwestern Sierras Pampeanas are considered to be highergrade equivalents of this formation (Toselli 1990; Willner 1990). The predominant lithology consists of turbiditic clastic sequences, corresponding to the channel, proximal, intermediate, and distal facies of a prograding submarine fan system (Je~ek 1990 and references therein). Limestones, dolomites and marbles (in the higher-grade

THE PAMPEAN OROGENY rocks) are interbedded in the uppermost level of the turbidite sequence (Porto et al. 1990), while alkaline lava flows typical of rift-related environments are interlayered with the clastic sequences (Rapela et al. 1992 and references therein). Palaeocurrent directions in the clastic sequences indicate provenance from the east and southeast, swinging to north-south in the most distal facies, suggesting that all these sediments were deposited along a passive continental margin on the western side of Gondwana (Je~ek 1990). The low-grade clastic facies have provided diverse and relatively abundant trace fossils, with more than 300 specimens of soft-body impressions corresponding to 23 ichnogenera, all indicating a late Precambrian (Vendian) to early Lower Cambrian (Tommotian) age (Durand 1996). The age of the high-level tight chevron-folding that affects the Puncoviscana Formation s.I. is therefore stratigraphically constrained between the minimum age for sedimentation (Tommotian, c. 530-534 Ma) and the age of the overlying Late Cambrian Mes6n Group (Rossi et al. 1992 and references therein). That leaves an important hiatus in the sedimentation, c. 530-505 Ma, during which the main deformation and metamorphism occurred and granitic intrusions were emplaced, together referred to as the Pampean Orogeny (e.g., Rossi et al. 1992; Durand 1996). Isotopic data are sparse but tend to confirm the above geological observations: U-Pb analyses of detrital zircons in the northern low-grade rocks indicate a maximum age of 530-560 Ma for deposition of the Puncoviscana Formation, with a Pb-loss episode at 527Ma (Lork et al. 1990). K - A r ages on micas from the same sector are 535-540Ma (Adams et al. 1990), while Cordani et al. (1990) reported whole-rock Rb-Sr isochrons on low-grade pelites of 5 3 8 + 2 3 M a (initial 87Sr/867Sr= 0.7126, M S W D = 0 . 7 ) and 520+ 14Ma (initial 87Sr/86Sr = 0.7148; MSWD = 0.4). For the highgrade equivalents of the Puncoviscana Formation in the northern Sierras Pampeanas, Rb-Sr dating of thin whole-rock slabs of a banded gneiss produced an age of 548 + 27 Ma, interpreted as the age of the first metamorphic event, with further evidence for overprinting at 451 ~ 11 Ma (Sierra de Ancasti, Bachmann et al. 1986). Detrital zircon populations in a highgrade rock from the Sierra de Quilmes show one fraction plotting close to concordia at 530Ma and an upper intercept age near 1022Ma (L. Heaman, pers. comm.). Finally, recent geochronological work on isolated basement outcrops in northern Chile indicate mid-Cambrian ages for the Belen schists (536-544Ma, Basei

211

et al. 1996) and the Mejillones metamorphic complex (524Ma, Lucassen et al. 1996), which are consistent with previous Cambrian ages obtained by Datum et al. 1990), strongly suggesting that the Pampean orogeny had an important role in the amalgamation of the proto-Andean basement. The new results in the SC demonstrate that the Early to Mid-Cambrian deformation is no longer restricted to NW Argentina, and the coeval granulite facies metamorphism and $2 deformation correspond to low structural levels of the Pampean orogeny. Considering the results of this work, the Pampean orogeny is regarded as a continental collision event that affected a large sector of the Andean proto-margin, from at least 17~ (Puncoviscana Formation outcrops) to 33~ in the southernmost Sierras Pampeanas. The protoliths of the metamorphic sequence of the SC are dominated by passive margin turbidite deposits and shelf facies (marbles) to the east (Kraemer et al. 1995), which cannot be distinguished from the typical protoliths of the Puncoviscana Formation and its metamorphic equivalents in the northwestern Sierras Pampeanas. The passive margin sequence started during the Neoproterozoic with the opening of a marine rift basin, flanked in the northern sector by the Arequipa-Antofalla massif to the west and the Amazonian craton to the east (Fig. 17). It is worth noting that the deep marine facies of the Puncoviscana Formation also extended northwards only as far as approximately 20~ (Durand & Acefiolaza 1990), whereas shallow marine carbonate facies of the same age extend farther north in Bolivia and western Brazil (Fig. 17, Tucavaca Group and Tamengo Formation, Durand 1996). An aulacogenic origin has been proposed for this northern area, with a possible triple-junction separating the Arequipa, Amazonian and Rio Apas cratonic massifs (Litherland et al. 1986; Avila Salinas 1992) (see Fig. 17). A common crustal history for the Arequipa-Antofalla and Amazonian cratons is also indicated by their overlapping Pb-isotope compositions (Tosdal 1996), which imply that the Arequipa-Antofalla craton is not an allochthonous terrane accreted in later stages to the Andean proto-margin.

T e c t o n i c e v o l u t i o n d u r i n g the E a r l y to M i d - C a m b r i a n

interval

An important change occurred in the Puncoviscana Formation south of 25~ where the deep marine basin was flanked to the east by the Rio de la Plata craton but to the west was open to the

212

C. W. RAPELA E T AL.

ocean (Durand 1996), namely the 'Puncoviscan' ocean. Although the passive margin sequence of the SC is considered equivalent to that in the Puncoviscana Formation s.l., there are other lithologies affected by the M2-$2 event that are also characteristic of the southern sector: (i) Thin amphibolite bodies interbedded with the metasedimentary rocks have geochemical and isotopic MORB signatures (Fig. 4), suggesting that ocean crust was involved together with marine sediments. The ultramafic rocks and associated amphibolites of the western side of the SC also exhibit MORB geochemical fingerprints (Escayola et al. 1996). It is not clear whether this ophiolite sequence was affected by the M2-$2 event; it is often associated in the field with ductile shear zones (Bonalumi & Gigena 1987) that slightly post-dated the metamorphic peak. (ii) The G la metaluminous group of granitoids of the SC is considered to be the southward extension of the plutonic-volcanic belt located along the easternmost Sierras Pampeanas (Lira et al. 1996), sharing its calc-alkaline and volcanic arc signatures. The U-Pb age of 530zL3Ma deduced for this suite suggests that intrusion was partially coeval with, or closely followed, sedimentation of the later units of the Puncoviscana Formation (Tommotian age; 530-534 Ma, scale of Gradstein & Ogg 1996). In turn, emplacement of the suite preceded the Early Cambrian M2 metamorphic peak by no more than about 10 Ma. A geodynamic scenario for the Neoproterozoic to Mid-Cambrian interval may be inferred from the above characteristics and the succession of events outlined in Fig. 2. lntracratonic, marine rift-related sequences, indicate Late Neoproterozoic break-up of the supercontinent, which became an assemblage of cratons and small massifs, such as the Amazon, Arequipa-Antofalla and Rio Apas. Development of a Neoproterozoic-Tommotian (c. 530-550 Ma) passive margin sequence is represented by the Puncoviscana Formation, along the southwestern edge of Gondwana (Fig. 17). Northeastward subduction of oceanic crust started along the supercontinent margin during the Early Cambrian (c. 530 Ma). Destruction of the Puncoviscan passive margin led to a rapid build-up of a thick accretionary prism against the continental margin. Notably, initiation of subduction is favoured when the oceanic lithosphere younger than c. 20 Ma (Cloetingh et al. 1989). At this stage, a continental margin magmatic arc was established on the proximal parts of the prism, and probably on the

Proterozoic continental craton, although the latter relationship has not yet been demonstrated in the Eastern Sierras Pampeanas. The model assumes that metamorphism during subduction occurred only in the deeper part of the prism, which was affected from an early stage by seaward verging thrusts, and progressively migrated to the back-thrust zone. Collision with a continental block occurred during the late Early Cambrian-early MidCambrian. As pointed out above, the metamorphic P T - t path and associated strongly peraluminous magmatism clearly record regional crustal thickening resulting in pressures >8 kbar and T > 800~ A simple subduction model cannot explain these features, whereas collision with a continental block located to the west-southwest accounts for crustal thickening and the heat source to produce the various magma affinities. Underthrusting of the buoyant, exotic continental block caused seaward slicing of the accretionary complex, including the shelf deposits and the pre-collision granitoid suite. A recent analogue of a similar tectonic scenario is the Shimanto belt in Japan, where Cretaceous-Palaeogene formation of an accretionary prism was followed by Miocene collision with a continental micro-block (Charvet et al., 1990; Stein et al. 1994). The origin of the 'exotic' continental block that collided with Gondwana is more difficult to establish. However, an important point emerging from the evolution of the SC is that the Cambrian collision closely followed disintegration of the supercontinent (Fig. 2). This strongly suggests the 'exotic block' namely the Pampean terrane, was an integral part of the breakup scenario pertaining at the PrecambrianCambrian boundary. Inherited zircon from the metapelitic high-grade migmatites as well as from the S-type cordierite-bearing granites give ages between 600 and 1400Ma (Figs 12, 14), suggesting provenance from Meso-to-Neoproterozoic basement. Two main alternatives are briefly discussed below, but an in-depth discussion of the models is beyond the scope of this work and will be presented elsewhere. (1) Following the Neoproterozoic episode of supercontinent break-up, a continental block was rifted away from the Arequipa-Antofalla or the Rio de la Plata cratons. A ridge-crest jump followed by eastward subduction closed the newly formed Early Cambrian Puncoviscan ocean, leading to a collision of the rifted microcontinent (the Pampean terrane) with the Gondwana margin (Fig. 17). A variant of this model is that the Puncoviscan ocean resulted from the opening and subsequent closure of an

THE PAMPEAN OROGENY

~

~

+ + t + + ++-EarlyCambrian ~'~4+ ~5 3 0~M a )+.A M C - f (ca. ~. ~ + + + + + + + ~ =-~ ~ lt?

~-~_~---~ GONDWANA ,

/

Oceaniccrust

_ - ' ~ ~ ' ~ " ~ - ~ ~ ' ~ - :! ~ ~ L a t e / N eoprot..... ,c fo Earl~ M a r g a r i n e sequence's: ~.... U ~ T k- facies: .--7........ ~. Off-shelf 40oS turbidites of the 9g PuncoviscanaF.(s.I.) centre ~; ~'"l,~,' ~ " Shallowmarinefacies =#i

Fig. 17. Early Cambrian (c. 530 Ma) reconstruction of the proto-Andean margin of South America, showing distribution of passive margin sedimentary units (modified from Durand 1996) at the time when subduction of ocean crust started forming an accretionary prism and a calc-alkaline volcanic arc along the eastern edge of the Sierras Pampeanas. The subsequent closure of this ocean due to collision of a micro-continent (the Pampean terrane), marked the onset of the Pampean orogeny during early Middle Cambrian times. AAC, Arequipa-Antofalla craton; AMC, Amazon craton; RAC, Rio Apas craton; RPC, Rio de la Plata craton.

ensialic marginal basin, associated with the initiation of Early Cambrian subduction. However, a detailed study of the sedimentological and compositional characteristics of the clastic sediments of the Puncoviscana Formation clearly indicates a passive margin setting, with a continental source area located to the east (Je2ek 1990). Both variants of this model imply that the Pampean terrane should not be isotopically distinct from the Arequipa-Antofalla craton. In neither case is a genetic relation implied between the Pampean terrane and the Precordilleran terrane accreted to Gondwana at a later stage during the Lower Palaeozoic (see Astini this volume; Keller et al. this volume). (2) Palaeogeographical models advocating location of eastern Laurentia adjacent to southwestern Gondwana during the Late Neoproterozoic (e.g. Dalziel et al. 1994), might offer an alternative origin for the Pampean allochthon. Following a rift-drift transition along the southern Appalachian Mountains and the Puncoviscana conjugate margins, sequential rifting of two large Laurentian continental blocks separated by sectors of oceanic crust may have started during Early Cambrian time. The Pampean terrane that

213

collided with Gondwana during the Cambrian would have been the first of these rifted Grenvillian blocks. The Precordilleran terrane accreted to Gondwana during the Ordovician (see Astini this volume) is regarded as a later rifted Laurentian block in this model. It is important to note that Famatinian subduction along the Gondwana margin started after a quiescent period of about 30 Ma (c. 520-490 Ma) following the Pampean collision.

The Early Ordovician high-Na plutons." Famatinian inner cordilleran activity The Ordovician high-Na plutons of the Sierra de C6rdoba were intruded into the Pampean highgrade basement that underwent rapid uplift after the Early to Mid-Cambrian compressional events. Their ages indicate that they are coeval with the early granitoids of the Famatinian arc located farther west, and with the subduction episode that started at approximately 500Ma, close to the Cambrian-Ordovician boundary (see Pankhurst et al. this volume). However, their composition is distinct from that of the calcic suites of the Famatinian arc. The strong depletion in incompatible elements such as K, Cs, Rb and U, and enrichment in Sr, suggest derivation from a mafic source at pressures at which plagioclase was not a residual phase; the low Yb values of the trondhjemite-tonalite unit (Table 3) also point to residual garnet. Initial 87Sr/S6Sr ratios are lower than those of the Cambrian suites, but only that of San Agustin is low enough (0.7045, Table 7) to be compared with mantle values. Peraluminous high-Na and Sr granitoids are found in some continental-arc environments such as the Cretaceous to Tertiary batholiths of coastal Peru. They form an inner cordilleran belt above the thickened keel of the Andes, located 300km inboard of the present trench and main plutonic/volcanic arc (Petford & Atherton 1996). Their distinct composition compared to their coastal counterparts has been ascribed to high-pressure melting of newly underplated basalts (Petford & Atherton 1996). High-pressure (_>15 kbar) partial melting experiments on basalt have shown that the melt composition is primarily high-A1 trondhjemitetonalite-dacite, tonalite-dacite, leaving a restite assemblage of garnet + clinopyroxene-4- hornblende (Drummond et al. 1996 and references therein). The obvious tectonic environment with which to associate the high-Na plutons of the SC is the coeval Famatinian magmatic arc to the west, which was emplaced in the low-grade

214

C. W. RAPELA E T A L .

equivalent of the Pampean metamorphic basement. The preferred tectonic interpretation for the Ordovician high-Na plutons is therefore as an inner cordilleran, widely-spaced plutonic assemblage, emplaced in the overthickened, high-grade Pampean foreland. A TremadocLlanvirn back-arc basin (Saavedra et al. this volume) separated the inner and main arcs, from the Puna southwards as far as the Sierras de San Luis. In the northwestern Sierras Pampeanas, Ordovician plutons with relatively low initial STSr/86Sr ratios were also considered to be part of an inner cordilleran arc (Rapela et al. 1990). Peraluminous high-Na trondhjemite units are also present in that sector at 25~ (Sierra de Cachi, Galliski et al. 1990; Galliski & Miller, 1989); ); they have given precise U-Pb monazite ages between 466+1 and 4 8 1 + l M a (Lork et al. 1991). The intrusion of high-Na plutons in the Sierra de Cdrdoba located 800km to the south, suggests that inboard high-pressure equivalents of the main Early Ordovician arc were emplaced along the whole Pampean foreland.

Conclusions Our interpretation of the sequence of preCarboniferous events in the SC (Fig. 2) has several important constraints for the evolution of the early Andean proto-margin, examples of which are given below. (a) The episodic metamorphic evolution took place mainly in Early Cambrian and Early Ordovician times. Consequently, the SC are not the westernmost exposures of the Proterozoic Rio de la Plata craton, although Early Palaeozoic overprinting of early structures cannot be discounted everywhere. (b) As there are no crustal discontinuities between the SC and the northwestern Sierras Pampeanas, an obvious correlation between the two sectors indicates that the Early to MidCambrian Pampean orogeny extended for more than 800km along the proto-Andean margin (Fig. 17). (c) The protoliths affected by the high-grade metamorphism and penetrative deformation ($2, M2, Fig. 2) are dominated by (i) flysch-type clastic deposits, (ii) a carbonate component restricted to the eastern side of the Sierras and (iii) a calc-alkaline suite of granitoids of Early Cambrian age (Gla group). The protoliths of (i) and (ii) are interpreted as syn-subduction deposits coeval with (iii), formed during destruction of a passive margin sequence equivalent to the Neoproterozoic-Tommotian Puncoviscana Formation.

(d) The characteristics of the Early Cambrian regional metamorphism and associated igneous rocks in the SC are typical of those formed during crustal thickening related to collision of continental blocks. (e) The N-MORB signature of the interlayered basic rocks in the passive margin sequence and of obducted ophiolites (Escayola et al. 1996) at the end of the Pampean orogeny are compelling evidence of Early Cambrian oceanic lithosphere southwest of the margin of the Gondwana cratons. This interpretation is consistent with a closing of the Puncoviscan ocean due to subduction of oceanic crust, and development of a volcanic arc represented by G l a metaluminous, calc-alkaline group of granites. (f) The age of the G l a subduction-related magmatism is constrained to the Early Cambrian by consistent U-Pb ages obtained on abraded zircons from granitoids of this group. Intrusion of this group pre-dated the M2 metamorphic event. Is noteworthy that the Late Cambrian interval shows no evidence of orogenic activity or subduction magmatism in either the Sierras Pampeanas (Loske & Miller 1996) or the Puna, where the Mid-Cambrian extension lasted until the early Ordovician (Bahlburg, 1990 & Herv6 1997). This might be a severe constraint to both the microcontinental (e.g., Astini et al. 1995) and the Laurentia-Gondwana collision models (Dalla Salda et al. 1992a; Dalziel 1997), since they propose protracted Cambrian subduction at the Gondwana margin followed by collision in Ordovician times. (g) At this stage we suggest, as a working hypothesis, the collision of a hypothetical block rifted from the Arequipa or the Rio de la Plata cratons shortly after the supercontinent breakup. An alternative model involves the Early to Mid-Cambrian collision of a Grenvillian terrane rifted from Laurentia at an early stage (i.e. prior to the rifting of the Precordillera terrane). In either case, Nd model ages and zircon inheritance indicate that the major rock-forming event in the deep crust of the SC was Palaeo- or Mesoproterozoic (c. 1500-1700 Ma). (h) Late Cambrian to earliest Ordovician intrusion of high-Na trondhjemitic and granodioritic plutons affected by dextral wrench shear record a coeval inner-cordilleran counterparts of the Famatinian arc. The eastward polarity for Famatinian subduction is thereby confirmed. This work was carried out as part of contract CI1CT92-0088" from the Commission of the European Communities and was also supported by grant PIP No. 4148 from the Consejo Nacional de Investigaciones Cientificas y T~cnicas de la Repfiblica Argentina (CONICET). It is presented as a contribution to

THE PAMPEAN OROGENY IGCP Projects 345 (Andean Lithospheric Evolution) and 376 (Laurentia-Gondwana Connections), and constitutes NIGL Publication No. 240. We are grateful to C. G. Barnes and L. Heaman for helpful reviews.

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KRAEMER, P. E., ESCAYOLA,M. P. & MARTINO, R. D. 1995. Hip6tesis sobre la evoluci6n tect6nica neoproterozoica de las Sierras Pampeanas de C6rdoba (30~176 Revista de la Asociaci6n Geol6gica Argentina, 50, 47-59. KRETZ, R. 1983. Symbols for rock-forming minerals. American Mineralogist, 68, 277-279. LA ROCHE, H. 1992. Un homologue cationique du triangle Q-A-P (quartz-feldspath alcalin-plagioclase), figure majeure de la p6trologie des roches plutoniques. Comptes Rendus de l'Academie des Sciences, Paris, 315, Serie II, 1687-1693. LE MAITRE, R. W., BATEMAN,P., DUBEK, A., KELLER, J., LAMEYRE, J., LE BAS, M. J., SABINE, M. A., SCHMID, R., SORENSEN, H., STRECKEISEN, A., WOOLLEY, A. R. & ZANETTIN, B. (eds) 1989. A Classification of Igneous Rocks and Glossary of Terms. Recommendations .of the International Union of Geological Sciences Subcommission on the Systematics of Igneous Rocks. Blackwell, Oxford. LINARES, E. & GONZALEZ, R. R. 1990. Catdlogo de edades radim~tricas de la Rep~blica Argentina 1957-1987. Asociaci6n Geol6gica Argentina, Publicaci6n Especial Serie B (Didfictica y Complementafia), 19. LIRA, R., R., MILLONE, H. A., KIRSCHBAUM,A. M. & MORENO, R. S. 1996. Calc-alkaline arc granitoid activity in the Sierra Norte-Ambargasta Ranges, central Argentina. Journal of South American Earth Sciences, 10, 157-177. LITHERLAND, M., ANNELLS, R. N., APPLETON, J. n . et al. 1986. The geology and mineral resources of the Bolivian Precambrian shield. British Geological Survey, Overseas Memoir, 9. LOSKE, W. P. & MILLER, H. 1996. Late Proterozoic and Early Paleozoic isotope data for geological events in W and NW Argentina. XIII Congreso Geol6gico Argentino y III Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, V, 529. LORK, A., MILLER, H., KRAMM, U. & GRAUERT, B. 1990. Sistem/ttica U-Pb de circones detriticos de la Formaci6n Puncoviscana y su signficado para la edad mfixima de sedimentaci6n en la Sierra de Cachi (provincia de Salta, Argentina). In: ACElqOLAZA, et al. q.v., 199-208. , , & 1991. U-Pb investigations of monazite and polyphase zircon: implication for age and petrogenesis of trondhjemites of the southern Cordillera Oriental, NW Argentina. 6 ~ Congreso Geol6gico Chileno, Vifia del Mar, Actas, I, 398-901. LUCASSEN, F., WILKE, H. G., VIRAMONTE, J., BECCHIO, R., FRANZ, G., LABER, A., WEMMER, K. & VROON, P. 1996. The Paleozoic basement of the central Andes (18~176 A metamorphic view. Troisieme Symposium International sur la GOod# namique Andine, Saint Malo, France, R6sumes 6tendus, ORSTOM 6ditions, Paris, 779-782. LUDWIG, K. 1980. Calculation of uncertainties in U-Pb isotope data. Earth and Planetary Science Letters, 46, 212-220. MARTINO, R, KRAEMER,P., ESCAYOLA,M., GIAMBASTIANAI, M. & ARNOSIO, M. 1995. Transecta de las

THE PAMPEAN OROGENY Sierras Pampeanas de Cdrdoba a los 32~ Revista de la Asociaci6n Geoldgica Argentina, 50, 60-77. MUIR, R. J., IRELAND, T. R., WEAVER, S. D. & BRADSHAW, J. D. 1996. Geochronology of Paleozoic magmatism in New Zealand and correlations with eastern Australia, Tasmania and Antarctica. Chemical Geology, 127, 191-210. OTAMENDI, J. E. & RABBIA,O. 1996. Petrology of highgrade gneisses from Macizo Rio Santa Rosa: evidence of decompression in the Eastern Sierras Pampeanas, Argentina. XIII Congreso Geol6gico Argentino y III Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, V, 527. PACES, J. B. & MILLER, J. D. 1993. Precise U-Pb ages of Duluth Complex and related mafic intrusions, northeastern Minnesota: Geochronological insights to physical, petrogenetic, paleomagnetic, and tectonomagmatic process associated with the 1.1 Ga Midcontinent Rift System. Journal of Geophysical Research, 98, 13 997-14013. PANKHURST, R. J. & RAPELA, C. W. 1994. Production of Jurassic rhyolite by anatexis of the lower crust of Patagonia. Earth and Planetary Science Letters, 134, 9-23. -& - - 1 9 9 8 . Introduction. This volume. , SAAVEDRA,J., BALDO, E., DAHLQUIST,J., PASCUA, I. & FANNING, C. M. 1998. The Famatinian magnetic arc in the central Sierras Pampeanas. This volume. PEARCE, J. A., HARRIS, N. B. W. & TINDLE, A. G. 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of Petrology, 25, 956-983. PI~REZ, M. B., RAPELA, C. W. & BALDO, E. G. 1996. Geologla de los granitoides del sector septentrional de la Sierra Chica de Cdrdoba. XIII Congreso Geol6gico Argentino y III Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas V, 493-506. PETFORD, N. & ATHERTON, M. 1996. Na-rich partial melts from newly underplated basaltic crust: the Cordillera Blanca Batholith, Peru. Journal of Petrology, 37, 1491-1521. PORTO, J. C., FERNANDEZ, R. I. & CARRION, M. H. 1990. Calizas y dolomias de la Formaci6n Puncoviscana, S. L. In: ACENOLAZA,et al. q.v., 37-52. RAPELA, C. W., COIRA, B., TOSELLI,A. & SAAVEDRA,J. 1992. The Lower Paleozoic magmatism of southwestern Gondwana and the evolution of the Famatinian orogene. International Geology Review, 34, 1081-1142. , PANKHURST, R. J., KIRSCHBAUM, A. & BALDO, E. G. 1991. Facies intrusivas de edad carbonifera en el Batolito de Achala: Evidencia de una anatexis regional en las Sierras Pampeanas? 6 ~ Congreso Geol6gico Chileno, Vifia del Mar, Actas, I, 40-43. - - , TOSELLI, A., HEAMAN, L. & SAAVEDRA,J. 1990. Granite plutonism of the Sierras Pampeanas: An inner cordilleran Paleozoic arc in the southern Andes. In: KAY, S. M. & RAPELA, C. W. (eds) Plutonism from Antarctica to Alaska. Geological Society of America Special Papers, 241, 77-90. RossI, J. N., TOSELLI, A. J. & DURAND, F. R. 1992. Metamorfismo de baja presi6n, su relaci6n con el

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Laurentia-Gondwana collision: the origin of the Famatinian-Appalachian Orogenic Belt (a review) L U I S H. D A L L A CARLOS

S A L D A 1, M O N I C A

A. C I N G O L A N I

G. L O P E Z

1 & RICARDO

D E L U C H I 2,

V A R E L A 1.

1 Centro de Investigaciones Geol6gicas, Universidad Nacional de La P I a t a - C O N I C E T , Calle 1 No. 644, 1900 La Plata, Argentina 2 Centro de Investigaciones en Recursos Geoldgicos-CONICET, Ramirez de Velazco 847, 1414 Buenos Aires, Argentina Abstract: Evidence from Cambrian to Devonian palaeogeography, the Laurentian prov-

enance of Precordillera, and comparison of Lower Palaeozoic tectonic, metamorphic, and magmatic events, suggests a common collisional origin for the Taconian Appalachians and the Famatinian (Puna, Sierras Pampeanas, North Patagonian Massif) orogenic belt. Detached from Eastern Laurentia following collision, the Occidentalia terrane (west of the Famatinian orogenic belt) is partly covered by the Precordillera carbonate bank, which contains syn-collisiona1460 Ma K-bentonites, related to the Puna and Famatinian arcs. In both the Appalachian and the Famatinian belts, a mostly Cambrian pre-collisional subduction regime developed at the western margin of Gondwana, following the opening of Iapetus. Laurentia Gondwana collision produced deep thrusting and crustal melting. As in the Blue Ridge of the Appalachian Mountains, oceanic and mantle rocks, and slices of Grenvillian basement were emplaced along the western side of the Famatinian belt. This hypothesis gives continuity to a paired Taconian Famatinian tectonic setting: a western Lower Palaeozoic platform overlying Grenvillian basement, and an eastern complex of a Lower-Middle Palaeozoic igneous-metamorphic belt with Proterozoic inliers. Prior to the idea of a South West LaurentiaEast Antarctica connection (Moores 1991; Dalziel 1991), palaeomagnetic reconstructions only showed a blurred image of the South American plate (SAM): as seen from the distant northern hemisphere. Dalziel (1991) drew attention to possible N A M (North America)-SAM interactions, and it was suggested that Taconian continent-continent collision could explain southern Iapetus evolution, including the Famatinian orogenic belt, the Occidentalia terrane, and Precordillera origins (Dalla Salda et al. 1992a, b; Dalziel et al. 1994). Recent palaeogeographic and palaeomagnetic work allows more than one possible position of N A M with respect to SAM in the Lower Palaeozoic. These include an archetypal one in which, at the moment of the Taconian orogeny, SAM was separated from N A M by c. 4000 km of oceanic crust (e.g. Torsvik et al. 1995), and an alternative in which these continents drifted apart during Cambrian times, but interacted during Taconian times (Dalziel et al. 1994; Dalziel 1997). In the southern Iapetus region, the second alternative is supported by detailed geological evidence, such as data from the northern Chile Palaeozoic basins (see Bahlburg 1993), the presence of the Precordilleran Lower Palaeozoic carbonate bank covering a Grenvillian basement, and Lower to Middle Palaeozoic

igneous-metamorphic belts in both N A M and SAM. Thus, Mid-Ordovician interaction can explain much of the geology of eastern N A M and western SAM, as well the origin of the Precordillera. MacNicaill et al. (1997) renewed the archetypal scenario, considering Ordovician palaeomagnetic and faunal data in an attempt to reconstruct evolution of the Iapetus Ocean. They preferred a palaeogeographical scenario with a wide southern Iapetus Ocean during Taconian times, in which eastern N A M faced three arcs, and the Precordillera terrane was drifting from N A M to SAM, as proposed by Astini et al. (1995) and Thomas & Astini (1996) to explain differences in faunal provenance in the Precordillera. MacNicaill et al. (1997) favoured the 'funeral ship' hypothesis for the 'Precordillera terrane', which was interpreted as a minor terrane detached from the Ouachita embayment during Cambrian differential transform faulting (Thomas & Astini 1996). Another recent palaeogeographic reconstruction, based on palaeomagnetic data and magmatic arc petrology, placed part of the Famatinian orogen (the Eastern Puna and Famatina magmatic arcs) in the southern Iapetus Ocean, close to the Gondwana margin and to some of the Lower-Middle Ordovician central mobile belt terranes of the northern

DALLA SALDA, L. H., LOPEZ DE LUCHI, M. G., CINGOLANI, C. A. et al. 1998. Laurentia Gondwana collision: the origin of the Famatinian-Appalachian Orogenic Belt (a review). In: PANKHURST,R. J. & RAPELA,C. W. (eds) The Prata-Andean Margin of Gandwana. Geological Society, London, Special Publications, 142, 219 234.

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Appalachians (Conti et al. 1996). In this reconstruction, the accretion of the Eastern Puna-Famatinian arc may have resulted from the closure of the Iapetus ocean during MidOrdovician times. A variant of the collisional model considers the Precordillera as a 'Malvinas Plateau-like' extension of NAM (Dalziel 1997). Evolution of the Palaeozoic rocks of NW Argentina (Coira et al. 1982), and of a narrow belt of Eastern Sierras Pampeanas (Ramos 1988), has also been interpreted in terms of continent-continent collisions. Dalziel & Forsythe (1984) suggested a possible southern extension of the Arequipa Massif to have collided with the Puna. Ramos et al. (1986) proposed an allochthonous microcontinent (Chilenia), which collided with western South America during Devonian times. Dalla Salda et al. (1990, 1992a, b) proposed a collisional event for the simultaneous evolution of the North Patagonian Massif and the Sierras Pampeanas during the 510-360Ma Famatinian Cycle (Acefiolaza & Toselli 1976). Similarities between the Lower Palaeozoic carbonate lithofacies and faunas of both the Appalachians and the Precordillera of SAM (e.g., Ramos et al. 1986; Astini et al. 1995) suggest that the Precordillera was an allochthonous terrane accreted to SAM during the Palaeozoic (Ramos et al. 1986). Dalla Salda et al. (1992a, b) considered the evolution of the Proterozoic to mid-Palaeozoic rocks of southwestern SAM as related to a Taconian-Famatinian collision of a single Occidentalia terrane, comprising the Precordillera, the Puna, the Famatinian magmatic arc, the Sierras Pampeanas, and the North Patagonian Massif. They suggested reversion to the earlier use of Taconian in place of Ocloyic (Turner & M6ndez 1975) for the name of this orogeny in SAM. The Taconian Appalachian orogen was also related to a collision; but it there is disagreement over the identity of the colliding terrane - minor continents, island arcs (Rodgers 1982; McKerrow et al. 1991; Van der Pluijm et al. 1995) or Gondwana (Dalla Salda et al. 1992a, b).

The Occidentalia terrane The proposed Occidentalia terrane is a long, relatively narrow, crustal block composed of Neoproterozoic rocks, trending nearly N-S west of the Famatinian belt. It crops out along the Andes as minor remnants of mostly Grenvilleage crystalline rocks, from the Arica re-entrant of Peru to Patagonia (Fig. 1), partially covered by the Cambrian to early Ordovician carbonate

platform of the Precordillera. The eastern limit of Occidentalia is defined by a c. 100kin-wide suture zone within the Western Sierras Pampeanas low-grade metamorphic units. It bears poorly dated basi~ultramafic rocks, considered as formed in the Taconian orogeny. Baeza & Pichowiak (1988) concluded that: ~the northern Chilean crust evolution can be traced to the mid-Proterozoic (1250 Ma)'. U-Pb zircon ages varying from 1171 to 1080Ma have been published recently for Middle Proterozoic rocks in the Arequipa-Antofalla craton of western Bolivia and northern Chile (Tosdal 1996). Caminos (1979) characterized the Western Sierras Pampeanas (eastern Occidentalia) as a mainly low-grade metamorphic belt, in which granitoids are absent, and mafic-ultramafic rocks are common. It includes metamorphic rocks of still uncertain age, and inliers of older basement. Rb-Sr dating of high-grade schists and gneisses from the core of Sierra de Pie de Palo anticlinorium (immediately east of the Precordillera) gave the first hint of possible Grenville rocks east of the Andes (Varela & Dalla Salda 1992), now confirmed by recent data (1021+ 12Ma, Pankhurst & Rapela this volume). Preliminary 2~176 ages of 1091 938Ma have been reported for magmatic and metamorphic zircons from a variety of metamorphic rocks in the same area (McDonough et al. 1993). A suite of granitic orthogneisses from the Sierra de Umango have been dated at 1030+ 30 Ma, and gneisses and granulites from Sierra de Valle F6rtil at 963 + 8 6 M a (Varela et al. 1996). Precordilleran basement xenoliths in Miocene volcanic rocks yielded 1102 + 6 Ma zircon ages from mafic rocks, and 1118 + 54 Ma from zircons of acidic rocks (Kay et al. 1996). Detrital zircons in Silurian sediments also present Grenvillian ages (Loske 1992). The recognition of Caradocian basic magmatism (mafic lavas, dykes and sills) along the western margin of the Precordillera (Kay et al. 1984) led to the proposal of a Chilenia terrane accreted to the western margin of southwestern SAM during Devonian times (Ramos et al. 1986). Recently, Ramos & Basel (1997) presented a Grenvillian U-Pb age (1069 Ma) from the Cordillera Frontal crystalline basement, also to the west of the Precordillera, and which they considered comparable to their Cuyania terrane (the Precordillera and surrounding basement rocks see Ramos et al. this volume). They concluded either that the gneisses belong to the Chilenia basement, and Cuyania and Chilenia have a common age and origin, or that there is a tectonic slice of Cuyania along the eastern border of the Cordillera Frontal.

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Fig. 1. Sketch map of the Famatinian orogenic belt (shaded) of SAM, showing the main geological features referred to in text. The Occidentalia terrane, comprising the basement of the Puna, the Sierras Pampeanas, and the North Patagonian Massif, is considered to be a slice of eastern NAM.

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South of Cordillera Frontal, in the San Rafael area (Ponon Trehue, Mendoza province), zircons from metamorphic rocks underlying Ordovician sedimentary rocks, are reported as having U-Pb Grenvillian ages (Astini et al. 1996). These outcrops are also included in Cuyania by Ramos et al. (this volume). Further east, in the San Luis region, (southwestern Sierras Pampeanas) some gneisses showed muscovite 4~ 1200800Ma ages (Gardini et al. 1996). In synthesis, we suggest that all the above mentioned data (see Fig. 2) give support to the

idea of the Occidentalia terrane as a major Proterozoic crustal belt in western SAM.

The Famatinian Orogen in the Puna and the Sierras Pampeanas Lying east of Occidentalia (Fig. 1), the Famatinian orogen is a complex metamorphic-igneous belt that evolved from Neoproterozoic to Devonian. It includes the Puna, the Famatina magmatic arcs, the Central-Eastern Sierras Pampeanas, where Acefiolaza & Toselli (1976) distinguished Pampean (590-530 Ma) and Famatinian (510-360Ma) cycles, and the North Patagonian Massif (Dalla Salda et al. 1992a). The Pampean

Fig. 2. The Grenvillian rocks of the Occidentalia terrane in central-west Argentina. Note the continuity of Grenvillian basement beyond the Precordillera Lower Palaeozoic carbonate bank. The Iapetus transition zone is shown east of Occidentalia, within the metamorphic rocks of the Western Sierras Pampeanas (suggested equivalents to the NAM Blue Ridge belt); it contains Grenvillian rocks as inliers. To the east lie the igneous-metamorphic rocks complexes of the Famatinian orogenic belt.

cycle

The Pampean cycle encompasses sedimentary, magmatic and metamorphic events that took place between the Neoproterozoic and the Lower Cambrian. It includes the first record of a post-rift magmatism along the western border of Gondwana, developed in ancient continental crust. Dalla Salda (1987) related it to an active margin episode, facing a pre-Pacific margin of Gondwana during the Cambrian. The protoliths of the metamorphic rocks that make up the Sierras Pampeanas, such as the Neoproterozoic to lowermost Cambrian Puncoviscana Formation that bears Oldhamia ichnofossils (Acefiolaza & Durand 1986), could be rift-related. Scattered meta-calcareous units in Cumbres Calchaquies, the Sierra de Ancasti, and the Sierra de Fiambal/t may also represent remnants of this margin. Evidence of Pampean magmatism is documented in the Eastern Sierras Pampeanas, where Rapela el al. (1990) defined a Late Precambrian to Cambrian G1 group of gabbro, tonalite and metamorphosed norite, with tholeiitic to island arc signatures and associated with a low-grade metamorphic event. Similar rocks have been described in the Sierra de Fiambal/t (DeBari 1994) and in the Sierra de Ancasti. A pre-530 Ma metamorphic event is deduced from K-Ar and Rb-Sr ages of gneisses, amphibolites and schists from the Sierras de Cdrdoba (Linares & GonzAlez 1990). Lira et al. (1997) considered some of these rocks as representative of a precollisional volcanic arc setting, spatially related to an active continental margin. In contrast, the 520 Ma age for E1 Pildn granite (Rapela et al. this volume) indicates a deep crustal melting episode and implies a thickened crust, probably associated with ocean basin closure along the above mentioned eastern margin. The Iruyica

LAURENTIA-GONDWANA COLLISION unconformity (c. 530Ma), over earlier passive margin sequences, may be interpreted as representing the end of the Pampean cycle at the autochthonous SAM margin.

The Famatinian

o'cle

The Famatinian cycle is an extensive period of metamorphism, igneous activity and older crust reworking, that extends from Late Cambrian to Devonian. Part of the granitoids of this cycle are equivalent to the 500-360Ma G2 granitoids proposed by Rapela et al. (1990, 1992). Magmatic rocks of broadly Ordovician age are widespread in NW Argentina (Lork & Bahlburg 1993). Basic and acidic lavas and associated volcaniclastic turbidites exposed in the eastern Puna extend southwards to the Famatina magmatic arc. Intrusion of 476-467Ma granitoids into folded Ordovician sedimentary rocks has been described (Bahlburg 1990). Geochemical data have been interpreted as indicating either a magmatic arc signature (Ramos 1988; Rapela et al. 1992), or a collisional one (Lork & Bahlburg 1993). The latter idea seems more feasible taking into account the evolution of tectonic basins (Allmendinger et al. 1982; Moya 1988). In the western Puna, Ordovician magmatism yields ages ranging from 487 to 441Ma, (Palma et al. 1986). Remnants of an older metamorphic basement, with poorly constrained ages, separate these western and eastern Puna magmatic belts. Three main groups of granitoids are distinguished on the basis of geochemical features and their relation to D2 and D3 deformational events in the Sierras Pampeanas (Ldpez de Luchi & Dalla Salda 1995): a pre-Famatinian group, a Famatinian group, and a post-Famatinian group. (1) The pre-Famatinian group (Late Cambrian-Early Ordovician, 520-460 Ma) comprises roughly concordant plutons with extensive mylonitization and recrystallization. These granitoids are calc-alkaline, metaluminous to peraluminous, biotite or hornblende-biotite tonalites to granodiorites, such as those from Sierra de San Luis (Sato & Llambias 1994). Major and trace element ratios also point to a volcanic arc signature. 87Sr/86Sri varies between 0.7043 and 0.7097, and primitive Rb-poor magmas evolve with a moderate increase in both Rb and Nb + Y (Fig. 3). (2) The Famatinian granitoid group (460420Ma) includes plutons concordant with the regional foliation. They are essentially calcalkaline to alkali-calcic, peraluminous biotitemuscovite monzo- to syeno-granites, either

223

associated with extensive granitic migmatization or intruded into the pre-Famatinian granitoids. In their restricted composition, peraluminous character, and LIL-enrichment, they are akin to collisional types. 878r/86Sri varies between 0.706 and 0.721 and there is Rb enrichment (Fig. 3). The peraluminous N-S elongated stocks of the Famatinian magmatic arc, i.e. Los Ramblones, Las Trancas, Loma Colorada; and the 430Ma Potrerillos pluton in the Sierra de Paimfin can also be included in this group. (3) The post-Famatinian granitoid group (420-330Ma) is characterized by two main fault-controlled batholiths. A N N E fault-controlled set of plutons show peraluminous to slightly peralkaline monzogranitic to syenogranitic compositions and contact aureoles. In the central Sierras Pampeanas they present, they bear magmatic andalusite, cordierite and even sillimanite (Rapela et al. 1996; L6pez de Luchi & Dalla Salda 1997). They are MgO-poor crustal granites with collisional to within-plate trace element signatures and high 87Sr/S6Sri (0.710 0.7197); shallow emplacement was proposed for these rocks (Rapela et al. 1996). Another set, of WNW-ESE fracture-controlled plutons (380330Ma), is mainly of metaluminous to mildly peraluminous, Mg-rich monzogranites to syenogranites with a mixed mantle-crustal signature. S7Sr/86Sri varies between 0.7057 and 0.7077, and trace element patterns point to a volcanic arc and within-plate signature (Fig. 3). Major and trace element patterns of the associated enclaves and dykes indicate mass and advective transfer from subduction-enriched mantle melts during the ascent and emplacement process (L6pez de Luchi 1996). Rapela et al. (1990) indicate that some final mid-Palaeozoic granitoids have alkaline affinities. Medium pressure high-temperature metamorphism is characteristic of the Taconian collision and some of the epizonal late- to post-Famatinian granites in the Sierras Pampeanas show a cordierite-andalusite aureole. Higher pressure conditions developed locally, in some cases representing unequilibrated pre-collisional metamorphism. A complex deformation affected the Famatinian belt, in which interference patterns and distinctive changes of structural trends are common (Dalla Salda 1987). D2 includes NNW-trending foliation, folding of S1 schistosity, and mylonitization developed during a ductile deformation episode associated with metamorphism and migmatization. D3, NNEtrending, is a higher crustal level episode associated with retrogressive metamorphism and mylonitization. Different crustal levels were

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Fig. 3. Geochemistry of selected samples of Famatinian cycle granitoids from the Sierras Pampeanas. Insets show TAS, Shand index, AFM, and Rb-Y + Nb diagrams. Key to fields: solid line, pre-Famatinian granitoids; dashed line, Famatinian granitoids; dotted line, post-Famatinian granitoids.

brought into contact by westward-vergent thrusting in the Puna (Allmendinger et al. 1982) and the Sierras Pampeanas (Toselli et al. 1992). Willner (1990) described shear belts in the central Sierras Pampeanas as representing a major Lower Palaeozoic crustal thickening event.

Neoproterozoic to Lower Palaeozoic Patagonia The North Patagonian Massif (Fig. 4), about the size of France, contains the largest and most important exposures of pre-Jurassic rocks in Patagonia. Ramos (1988) interpreted Patagonia as an allochthonous terrane accreted to the southern tip of South America in Upper Palaeozoic times. Consideration of its Neoproterozoic to Middle Palaeozoic igneous-metamorphic complexes, granitoid evolution, the

origin and trend of D2 foliation, and palaeomagnetic reconstruction, lead us to suggest that Patagonia, like the Sierras Pampeanas-Puna, also evolved as a result of the Taconian Laurentia-Gondwana collision during Ordoviclan times. The Proterozoic Colohuincul complex (Fig. 4) crops out from 39~ to 42~ along the eastern slopes of the Andes (Dalla Salda et al. 1990). Schists, quartz-phyllites, sillimanite gneisses, migmatites, amphibolites and minor granulites characterize this unit. A few small Neoproterozoic to Early Cambrian granitoids with a magmatic arc signature intrude the eastern edge of the complex, mainly biotite and biotiteamphibole tonalites and granodiorites conformable with the regional foliation (555+20Ma, 6 2 0 + 3 0 M a , Dalla Salda et al. 1991). The metamorphic rocks have yielded an Rb-Sr age of up to 860 4-30Ma and two Sm-Nd model

LAURENTIA-GONDWANA COLLISION

225

Fig. 4. The pre-Andean geology of the North Patagonian Massif.

ages of c. 1250Ma (Dalla Salda et al. 1991). It was suggested that the protolith of the complex was a remnant of a Proterozoic continental slab with Grenvillian affinities, the Colohuincul terrane (Dalla Salda et al. 1990). The Rio Chico complex, a wider central belt that runs east of the Colohuincul complex (Fig. 4), consists of a metamorphic unit (the Cushamen unit) and an Ordovician to Silurian batholithic-scale granitic suite (the Mamil Choique granitoids, Dalla Salda et al. 1994). The Cushamen unit experienced high-temperature medium-pressure metamorphism, during which tonalitic and granodioritic migmatites were formed, in association with deep crustal mylonites. The main foliation trend (NNW-SSE) was assigned to an Ordovician regional tectonothermal event (Dalla Salda et al. 1994). The Mamil Choique granitoids are biotite-epidote and two-mica granodiorites and monzogranites, which show a similar foliation trend to the metamorphic rocks. Their restricted acidic compositions and peraluminous compositions with high normative corundum indicate crustal sources. The easternmost region contains the infracrustal Mina Gonzalito Yaminu6 gneissic complex (dated at c. 800 Ma by Linares et al. 1990, and c. 650Ma by Varela et al. 1997), and the supracrustal E1 Jaguelito-Nahuel Niyeu low-

grade metapelites (c. 600Ma, Linares et al. 1990). The latter unit is intruded by a few Lower to Middle Palaeozoic granites, such as the Rio Salado granodiorite dated at 476 + 4 Ma (U-Pb zircon, Varela et al. 1997). The two units constitute the E1 Jaguelito terrane, which was considered part of the autochthonous craton of southwest Gondwana by Dalla Salda et al. (1991); Chernicoff (1994) included its evolution in the Pampean cycle. Recent U-Pb zircon data yield an age of 526 :t: 15 Ma for the Mina Gonzalito gneisses, suggesting the possible extension of the Brazilian orogenic belt into eastern Patagonia (Varela et al. 1997). Devonian plutons (c. 370 Ma) emplaced in the Rio Chico complex are undeformed calc-alkaline, peraluminous to metaluminous, epizonal, mainly two-mica leuco- and monzo-granites. A subsequent extensional tectonic regime is indicated by a few c. 350Ma K - A r ages for alkaline dykes (Dalla Salda et al. 1991).

The Taconian Orogeny in the Appalachian Mountains Neoproterozoic-Early Palaeozoic rifting and opening of the Iapetus Ocean produced the continental margin of Laurentia, along which

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the Appalachian-Ouachita orogenic belt subsequently formed (Thomas 1991). Neoproterozoic units correspond to continental, near shore environments, bearing Ohldamia ichnofossils from New York to Nova Scotia (as in the Puncoviscana Formation in western SAM). Components of the Appalachian-Ouachita rift are diachronous: rift-related rocks are overstepped by postrift Lower Cambrian strata in the Appalachian Blue Ridge and by Upper Cambrian strata along the Ouachita margin (Ham et al. 1964). Taconian deformation of the Appalachian orogen is interpreted as resulting from continent and/or island arc collisions during the closure of one or more Neoproterozoic to Palaeozoic oceans (Rankin et al. 1989). The history of Neoproterozoic continental rifting followed by the development of an Early Palaeozoic continental margin along the western side of the orogen is taken as conclusive evidence that the Taconian orogeny was preceded by continental break-up and formation of the Iapetus Ocean. On the eastern side of the orogen, differences in the lithostratigraphy and deformation of the accreted terranes are the basis for various hypotheses concerning their origin and accretionary history. The eastern NAM margin faced open ocean in Early Cambrian times, when the transition from rifting occurred (Hatcher 1989). A passive margin carbonate bank became established on the platform. Clastic sedimentation was probably continuous from early Mid-Cambrian to early Mid-Ordovician, but after the interruption of the Hawke Bay event in Tommotian times sedimentation retreated towards the west. This major change is thought to represent the rift-drift transition and establishment of the Laurentian trailing-edge margin. The Blue Ridge basement axis related to the Taconian event (i.e. the Iapetus suture) divides western Grenville basement and its cover from the main Appalachian allochthonous terranes (Hatcher 1989). These are, from west to east, the Piedmont zone, the Carolina terrane (Charlotte and Carolina slate belts), and the RaleighKiokee belt. The Carolina-Avalonia Terrane is composed of a Neoproterozoic to Cambrian assemblage of metamorphosed felsic, intermediate, and mafic volcanic rocks, with associated plutons and is similar to the successions that characterize Eastern Sierras Pampeanas. Pre-Taconian metamorphism has been recognized (Kish et al. 1979). The Neoproterozoic-Cambrian rockforming event in the Charlotte belt was interpreted as taking place in a peri-Gondwanan setting (Hibbard & Samson 1995).

The Piedmont zone, separated from the Blue Ridge by the Brevard Fault, is composed of medium-to-high grade metasedimentary and metavolcanic rocks enclosing minor maficultramafic rocks. Two belts have been distinguished (Hibbard & Samson 1995). The eastern belt is dominated by meta-clastic rocks associated with subordinate amphibolitic and ultramafic rocks. Two suites of plutonic rocks have been recognized in it, a mainly Cambrian one and a younger Late Ordovician-Silurian one that is coeval with development of a major unconformity as a result of the Potomac orogeny. The western belt corresponds to the accretionary complex of the eastern belt. The Potomac orogeny involves west-directed thrusting and is recognized along the western margin of the Piedmont zone. Although the Potomac orogeny could be unrelated to the Iapetan cycle, it is remarkable that plutonism subsequent to the unconformity shows restricted acidic composition, high peraluminosity, and marked LIL-enrichment. Late Precambrian to Early Cambrian plutonism also occurred along two belts: anorogenic peralkaline to peraluminous granites in the Blue Ridge (Drake et al. 1989) and a 680-540Ma calc-alkaline suite in the Carolina slate belt (Sinha et al. 1989). The differences in their implied tectonic settings, coupled with the presence of Atlantic faunas in the Carolina slate belt, led Sinha et al. (1989) to infer a large ocean basin during the Cambrian period. Late Cambrian to Devonian magmatism is well documented in the central and southern Appalachians, with plutons distributed in the Blue Ridge, Inner Piedmont, Charlotte, and Carolina slate belts (Sinha et al. 1989). Late Cambrian-Early Ordovician plutonism (520490Ma), represented mainly by multiplydeformed and metamorphosed tonalite and granodiorites, also includes rare granites showing a calc-alkaline trend and continental affinities. The root zones of the Cambrian arc are present, as in the Famatinian orogenic belt. A Taconian group of rocks (460-440 Ma) is distributed along the Blue Ridge, Inner Piedmont, and Carolina Slate belts. In the Inner Piedmont, it is related to the Potomac orogeny (Hibbard & Samson 1995), syn-kinematic peraluminous granitoids being emplaced into the metamorphic rocks. Accessory mineral phases indicate partial melting of a metasedimentary source under high to moderate pressure. Silurian-Devonian magmatism (420-360Ma) is present in two belts, one includes the Blue Ridge and the western area of the Inner Piedmont (associated with a postTaconian melting episode) and the other in the Charlotte belt (Sinha et al. 1989). In the former,

LAURENTIA-GONDWANA COLLISION peraluminous trondhjemites, granodiorites, and granites are conspicuous, whereas in the Charlotte belt, metaluminous monzogranites are common; however, there are also syenites, diorites, and gabbros related to transtensional tectonics during Acadian strike-slip accretion of the Carolina slate belt (Sinha et al. 1989). Taconian metamorphism developed a medium-pressure facies series in the Piedmont and a low-pressure facies series in the central Appalachians: metamorphic grade increased eastwards. In the Northern Appalachians, deformation of the accretionary prism and subduction-zone metamorphism during the close of the Iapetus Ocean have been proposed (Drake et al. 1989). In some places, the region of high-pressure metamorphism is bounded by shear zones (Laird & Bothner 1986) that bear ultramafic pods. In the central Appalachians, metamorphic zones of variable grade have been described, leading to the concept of a collage of accreted terranes. Actually, the low-pressure metamorphism could be either Taconian or older (Crawford & Mark 1982), and would have occurred on the continental side of an island arc. Dennis & Wright (1997) indicate that, in the western Carolina terrane (Charlotte belt), the main metamorphism is pre-Taconian (around 580535 Ma) and that it could have taken place in peri-Gondwanan reahns. The higher-pressure Taconian metamorphism would have resulted from crustal thickening. In the Southern Appalachians, a discontinuity in the metamorphism has been clearly delineated between western and eastern walls of the Hayesville fault (the inferred Taconian suture zone) with a Taconian (480-440 Ma) overprint in the Blue Ridge. Available data suggest that Taconian deformation in the northern Appalachians resulted from attempted subduction of the assembled stack of thrust sheets of the Laurentia continental margin under an island arc (Drake et al. 1989, Hatcher 1989). Models for the central Appalachians are more obscure, and two island arcs have been recognized, whereas in central Virginia east-dipping subduction was inferred. In the Southern Appalachians an ocean basin, between Laurentia and the eastern areas of the Appalachians, was proposed by Drake et al. (1989). Recently, MacNicaill et al. (1997) have suggested that the Taconian episode originated from the collision of NAM with an island arc that had evolved from Arenigian times.

Evolution of the Famatinian Orogen In SAM the Upper Precambrian Puncoviscana Formation and associated alkaline rocks point

227

to Neoproterozoic break-up. During drift, in the trailing-edge of eastern NAM a carbonate bank formed, while in western SAM an active margin developed, in which a Cambrian to Early Ordovician pre-Taconian magmatic arc was formed (Fig. 5). In Early Ordovician times Iapetus stopped widening, the NAM platform drowned, and SAM approached NAM as the ocean was closed by subduction under the SAM active margin. East of the Iapetus ridge, the most reliable record of subduction beneath SAM margin seems to be the pre-Famatinian Cambrian granitoids emplaced in the Puncoviscana Formation. These were related to arc magmatism (Omarini et al. 1996), and evolved to produce the c. 520 Ma

Fig. 5. Continental reconstructions for Cambrian (c. 530 Ma) and Ordovician (c. 485 Ma) times, based on data from Dalziel 1992 & Dalziel et al. 1994. OT, Occidentalia terrane (including Precordillera) restored to Ouachita embayment of Laurentia. SP, South Pole; NAM, Laurentia; GON, Gondwana; BAL, Baltica poles.

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L. H. DALLA SALDA E T A L .

Tastil and Cafiani late tectonic granitoids. The (Varela & Dalla Salda 1992), as in the Blue Ridge same authors considered that they had similar region of the Appalachians (Hatcher 1972). Intra-Ordovician tectonic episodes took place ages and tectonic setting to the trondhjemites of Newfoundland. Rapela et al. (this volume) have in the Precordillera. The first was at the start recognized three Cambrian igneous events in the of the Llanvirn (Guandacol tectonic phase, Sierras de C6rdoba: Early Cambrian subduc- Furque & Cuerda 1979), which we suggest can tion-related high/medium-grade metaluminous be related with the Penobscot disturbance that gneisses, slightly younger, strongly peraluminous predates deposition of the Lower-Middle Ordogranites that represent crustal anatexis during the vician in the Appalachians (Skehan 1988). The post-M2 thermal peak, and Late Cambrian to second, between the Upper Caradoc and the earliest Ordovician peraluminous trondhjemites Ashgill (Villicum tectonic phase, Baldis & that indicate high-pressure melting of old primi- Bordonaro 1982), was assigned to the Taconian tive material. A Neoproterozoic to Cambrian event. At the end of the Arenig the Cambrianpre-collisional granitic event occurred also in the Early Ordovician carbonate platform changed North Patagonian Massif, representing a sub- to black shale sedimentation in an interior basin duction-related initial phase of Early Palaeozoic (Baldis et al. 1984), as has been observed in the Appalachians. convergence. In the western Precordillera a marginal intraA c. 450Ma active margin, with a progressively thicker crust towards the east, has been continental basin evolved during Llandeilloproposed for the Famatinian magmatic arc Caradoc times. This intra-continental rift (Dalla (Toselli et al. 1996); however, older ages for Salda et al. 1992b; Loske 1992; Dalziel et al. the more basic rocks have been reported (Perez 1994) with Caradocian ophiolites, was inter& Kawashita 1992). In Patagonia a few Neo- preted as part of a post-Taconian rifting during proterozoic to Cambrian granitoids in the which the Ouachita trough developed (Dalla Colohuincul terrane-Rio Chico complex tectonic Salda et al. 1992b). Lowe (1985) indicated that front could also indicate a pre-collisional the southern block bounding the Ouachita Andean-type phase: the low S7Sr/86Sri (0.7021- trough could have consisted of a distended and 0.7032, Dalla Salda et al. 1991) and composition deformed microcontinent of Proterozoic sialic suggest upper mantle sources, and magmatic arc basement, along the southern margin of which lay the continuation of the Neoproterozoic affinities. The main stage of the collision, which reached Appalachian margin of NAM. In northern Patagonia syn-collisional rocks a climax around 460Ma, resulted in the Famatinian-Appalachian Taconian orogenic are represented by the medium-pressure metabelt (see Fig. 5). In the Famatinian orogen morphic rocks, migmatites, and granitoids of the this episode produced new crust by the growth Rio Chico complex, a consequence of crustal of syn-collisional arcs (Puna and Sierra de thickening resulting from the Taconian collision (Dalla Salda et al. 1994). Famatina granitoids), arc accretion (La Huerta In the Sierras Pampeanas, following collision island arc, Vujovich 1993), and granitization. Regional thrusting of the lithosphere produced and during the clockwise motion of Laurentia a 'piling up' of the upper crust, which underwent around Gondwana, mega-shearing, retrograde a low- to medium-P-high-T metamorphism and metamorphism, and emplacement of late to migmatization. The Famatinian granites were post-Famatinian granitoids took place. The conemplaced synchronously with the main deforma- tinental crust was progressively thicker towards tional phase (D2, trending NNW to NW, Dalla the Gondwana craton (Toselli et al. 1996). In Patagonia, the Proterozoic continental Salda et al. 1992a, b). Neoproterozoic to Lower Cambrian rift- plates involved in the collision were the Colorelated sediments were involved in the Taconian huincul terrane (a remnant of eastern Laurentia?) metamorphism in SAM (Toselli 1990). Deep to the west, and the E1 Jaguelito terrane to the water sediments of the eastern edge of the east. The syn-collisional migmatites and graniCambrian platform, and rocks related to island toids of the Rio Chico complex are also relicts of and continental arcs were also affected by the that event. In the North Patagonian Massif uplift collision. At the western edge of the Famatinian began during Late Ordovician to Early Silurian belt (e.g., in the Pie de Palo and E1 Gigante times, possibly providing the sediments which anticlinoria) low-grade marbles can be inter- filled up the mid to Late Palaeozoic Cordillera de preted as a distal part of the Lower Palaeozoic la Costa accretionary prism to the west. Several Precordillera platform (Baldis et al. 1984) Rb-Sr and K-Ar Devonian ages (c. 370 Ma) in involved in the Taconian collision. These rocks northern Patagonia characterize regional overare associated with a Grenville metamorphic core heating and deformation during the Acadian

LAURENTIA GONDWANA COLLISION event, and Devonian granites exhibit a collisional signature (Dalla Salda et al. 1994). We stress that the main metamorphic and igneous activity of the Famatinian orogenic belt and the Early Palaeozoic evolution of the North Patagonian Massif are Taconian, and preserve similar foliation trends, which controlled mainly eastdipping thrust planes. This major geotectonic cycle constructed the core of northern Patagonia as well as the central area of Argentina.

Similarities: the case for a c o m m o n o r o g e n

Geochronological data indicate a common Grenvillian crystalline basement to the west (and within, in small inliers) of both the NAM Taconian and Famatinian igneous metamorphic belts. A common Lower Palaeozoic carbonate bank partially covers the Grenvillian crystalline basement, and in both cases lies to the west of the Ordovician (Taconian) orogenic belt. To the east, some relicts of the SAM autochthon appear to be present. The similar stratigraphy, faunal content, and Ordovician collision-related K-bentonite levels of both Lower Palaeozoic carbonate platforms are remarkable; the Puna and Famatinian mid-Ordovician granitic arcs may explain the provenance of the K-bentonites (Huff et al. 1995; Cingolani et al. 1997). Both Taconian and Famatinian belts are Early to Mid-Palaeozoic collisional orogenic belts with an Ordovician peak of metamorphic and igneous activity (e.g., Hatcher 1972; Rodgers 1982; Tucker & Robinson 1990; Dalla Salda et al. 1992b). A geological fit, which includes the Famatinian, the Appalachians, the Ouachita Mountains, the Marathon ranges, the El Solitario uplift, and the Caborca area, as well as the common Iapetus transition zone, the probable edge of the Taconian autochthonous region of SAM, and the Neoproterozoic cratonic regions, also supports the hypothesis of the continuity of the Appalachians into the Famatinian orogenic belt (Fig. 6). The trend of Taconian foliation in NAM (Hall & Roberts 1988) and SAM metamorphic belts is similar and controlled the main thrust planes, dipping mainly east. In both belts some of these planes affected both the lower crust and the upper mantle (Dalla Salda et al. 1992a, b; Haworth et al. 1988). In NAM there is continental crust on either side of the Iapetus suture. In the west this is the autochthonous Grenvillian of NAM but, to the east, indirect data (zircon ages, geophysics) indicate a different crust. If our collision hypothesis is correct, the latter could be part of the

229

SAM autochthon, left behind after Taconian collision. It should be remembered that a Gondwanan origin has been suggested for part of the Florida basement (Mueller et al. 1994) and other circum-Atlantic exotic terranes (e.g., the Carolina, Avalonian, and Cadomian terranes). Several narrow, irregular, basic ultrabasic belts of Neoproterozoic and Early Palaeozoic age exist within both orogenic belts. The suture of the Taconian collision, fairly well delineated in the Appalachians (e.g., Williams 1979; Zen 1983), is considered to be mid-Ordovician. In SAM, equivalent rocks representing remnants of ocean floor, upper mantle, and lower crust slices, emplaced in short and narrow belts and lenses, now delineate a discontinuous syn-collisional suture along the Sierras de Valle Fbrtil, VirorcoLas Aguilas, Maz, Ancasti and Fiambalfi, west of the Famatinian orogenic belt. In summary, the Piedmont and Carolina zones would be broadly equivalent to the Central and Eastern Sierras Pampeanas (with older rocks towards the east), and the Blue Ridge equivalent to the Western Sierras Pampeanas low-grade metamorphic belt. The Virgiliana orogeny seems to be roughly equivalent to the Iruyic event. In both the Appalachian and Famatinian belts, a protracted tectono-thermal history developed after the Taconian event, lasting until Middle Devonian times (Acadian phase in NAM, Chafiic phase in Argentina). In the Southern Appalachians a 'severe continental collision' occurred during Acadian times (Robinson et al. 1988), which can be correlated with the proposed Acadian transpressive collision of NAM and SAM (Dalziel et al. 1994).

Conclusions

The similarities of the Famatinian orogenic belt and the Taconian Appalachians, and the ambiguity of Cambrian to Ordovician palaeomagnetic reconstructions, allow us to suggest the hypothesis of a common Taconian N A M SAM collisional origin for both belts. It can explain the nearly synchronous evolution of both belts, the residual terrane of Occidentalia in SAM and a possible relic of SAM crust in the western Appalachians, the sudden southward disappearance of the Appalachians, and restriction of the Famatinian arc to the south of the Arequipa block. We propose that the Occidentalia terrane, and Proterozoic inliers in the western Famatinian belt, are equivalent to the Grenville basement in USA, that were detached from eastern NAM coeval with origin of the Ouachita Mountains.

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500 km ]

Fig. 6. Fit showing common features of both Taconian and Famatinian belts. Note the continuity of the autochthonous margin of Laurentia with the Occidentalia terrane, the position of the Precordillera in relation to the Ouachita-Sonora-Caborca region, the proposed extension of the Famatinian belt south of the Appalachians, the suggested common edge of SAM (Cambrian active margin), and Iapetus suture transition zone.

LAURENTIA-GONDWANA The Argentine Precordillera is roughly equivalent to the C a b o r c a - S o u t h e r n Appalachians L o w e r Palaeozoic carbonate bank. In this c o m m o n T a c o n i a n orogen, the Occidentalia terrane with its Precordilleran L o w e r Palaeozoic c a r b o n a t e b a n k m a y be considered equivalent to the western A p p a l a c h i a n basin, the western Sierras P a m p e a n a s to the Blue Ridge belt, the central and eastern areas of the F a m a t i n i a n orogenic belt (including the N o r t h P a t a g o n i a n Massif) to the P i e d m o n t region, K i n g M o u n tains, Charlotte, and Carolina belts of Eastern Laurentia. In N A M , east of the Iapetus suture, some a u t o c h t h o n o u s p r e - T a c o n i a n S A M crust might be present, as exotic terranes detached after collision. A c o m m o n N A M - S A M eastward subduction margin, seems also to fit both belts, it points to a m o r e acceptable Southern Iapetus Cycle. Sometime before the end of the Ordovician, Laurentia and South A m e r i c a drifted apart and, at this m o m e n t , Laurentia lost the s o u t h e r n m o s t third of the Appalachians. L a u r e n t i a n drift left behind Occidentalia, taking away a slice of the a u t o c h t h o n o u s S A M that is t o d a y still east of the Appalachians Iapetus suture. The work was supported by grants of CONICET (Consejo Nacional de Investigaciones Cientificas, Argentine) and La Plata University. Special thanks to Ian W.D. Dalziel, Robert Hatcher Jr, Eldridge Moores, A.R. Palmer, Francisco Herve, William Thomas, Victor Ramos, Susan Kay, Carlos Rapela, and Heinrich Balhburg for opportune comments. Reviews of this manuscript by Robert Pankhurst, Chris Powell, and Bryan Storey were much appreciated. This paper is dedicated to Dr Roberto Caminos, a pioneer in Sierras Pampeanas geology. Drafting was carried out by M. Campafia. This is a contribution to IGCP Project 345.

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Structural evolution of the Sierra de San Luis (Eastern Sierras Pampeanas, Argentina): implications for the Proto-Andean Margin of Gondwana W. V O N

G O S E N 1 & C. P R O Z Z I 2

1 Geological Institute, University o f Erlangen-Nfirnberg, Schlossgarten 5, D-91054 Erlangen, Germany 2 Departamento de Geologia, Universidad Nacional del Sur, San Juan 670, 8000 Bahia Blanca, Argentina Abstract: The Western Basement Complex of the Sierra de San Luis underwent a complex deformational and metamorphic history of presumed Proterozoic age. The Phyllite Group and equivalents consist of a thick clastic sequence, whose time of deposition is compared with that of the Late Precambrian-Early Cambrian Puncoviscana Formation. This is intruded by early Famatinian magmatic arc tonalites to granodiorites, whereas an eastern basement unit was affected by early Famatinian mafic/ultramafic intrusions, metamorphism, and deformation. This Early Ordovician plutonism is related to east-directed subduction beneath the western margin of the Pampean terrane. Subsequent compression, plutonism, and variable-grade metamorphism are considered to be Mid- to Late Ordovician in age. Post-Famatinian c. WNW-ESE contraction under greenschist facies conditions led to folding of the cover sediments and uplift of different metamorphic units along mylonite zones which record an oblique sinistral component in the western part of the sierra. Based on a few radiometric data, this stage is interpreted as Devonian in age. It is possible that the compressive event was the result of the collision with the Precordillera (Cuyania) terrane in the west: oblique kinematic displacements, together with some evidence from outside the Sierra de San Luis, can be interpreted as the effect of the indenting terrane.

In western Argentina, the Sierras Pampeanas basement complex consists of two contrasting u n i t s - the Western and Eastern Sierras Pampeanas (sensu Caminos 1973, 1979; Dalla Salda 1987), on the basis of differences in lithology as well as magmatic and metamorphic evolution. The Andean half-horst structure of the Sierra (Grande) de San Luis represents the southernmost block of the Eastern Sierras Pampeanas. Its magmatic, tectonic, and metamorphic history spans ?Precambrian to Early Palaeozoic time and was terminated prior to the deposition of Late Palaeozoic clastic sediments. The Sierra de San Luis lies to the east of the inferred eastern margin of the Early Palaeozoic Precordillera (Cuyania) terrane (V. A. Ramos et al. 1996) or 'Texas Plateau' of Dalziel (1997), which also comprise most parts of the Western Sierras Pampeanas. It has been argued that this terrane collided with the Eastern Sierras Pampeanas during mid-Ordovician times (e.g., Astini et al. 1995; V. A. Ramos et al. 1996). The timing of collision and location of the suture zone, however, are still matters of debate. The magmatic, tectonic, and metamorphic history of the Sierra de San Luis can give an insight into the geotectonic evolution of this segment of the convergent western margin of Gondwana during the Early Palaeozoic era.

Although the area has been the subject of several regional and exploration studies (e.g., Kilmurray & Dalla Salda 1977; Kilmurray 1982; Romero 1984; Dalla Salda 1987; Hack 1987; Delakowitz 1988; Ortiz Su~rez 1988; Ortiz Sufirez & Ramos 1990; Prozzi 1990; Brodtkorb 1991; Ortiz Sufirez et al. 1992; von Gosen & Prozzi 1996), the magmatic, structural, and metamorphic evolution of many parts of the mountain chain is poorly understood and little is known about the timing of the different events. It has been suggested that in the metamorphic rocks, early folding around c. E - W axes was related to the 'Pampean Cycle' and that this was followed by the dominant N N E - S S W trend of the 'Famatinian Cycle' (e.g., Criado Roque et al. 1981; Dalla Salda 1987). However, clear separation of the pre-Famatinian structural and metamorphic events from those of the Famatinian Cycle has been difficult to establish. In the present paper we give a brief synthesis of the evolution of the metamorphic complex of the Sierra de San Luis. We describe the different lithological units and summarise the magmatic, structural, and metamorphic events during the pre-Famatinian, Famatinian and post-Famatinian cycles. The term 'Famatinian cycle' was used by Acefiolaza & Toselli (1976), Acefiolaza et al. (1978), and Acefiolaza & Miller (1982), to

VON GOSEN, W. & PROZZI, C. 1998. Structural evolution of the Sierra de San Luis (Eastern Sierras Pampeanas, Argentina): implications for the Proto-Andean Margin of Gondwana. In: PANKHUP,ST, R. J. & RAPELA, C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 235-258.

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comprise the post-Mid-Cambrian to Devonian time interval. New radiometric data from the Sierra de San Luis (Sims et al. 1997), show that the term 'Famatinian' can be restricted to Ordovician events. The older and younger events are ascribed generally to 'pre-Famatinian' and 'postFamatinian' cycles, respectively. The relative chronological order of events, based on the first deformational fabrics observed, and published radiometric data, will lead us to a preliminary interpretation of the timing of the structural and metamorphic overprints (summarized in Fig. 10) and subsequently to some larger-scale implications for the tectonic evolution of this part of the proto-Andean margin of Gondwana. In the text, the subscripts 'x' and 'F' refer to pre-Famatinian and Famatinian deformational events or fabric elements, respectively.

Lithological units The Phyllite Group ('San Luis Formation' of Prozzi & Ramos 1988) is exposed in two N N E SSW trending strips within the southwestern part of the sierra (Fig. 1). The monotonous succession consists of alternating phyllites and quartzites and is interpreted as a turbiditic sequence (Prozzi 1990). In the clastic deposits only two meta-conglomerate horizons have been found (e.g., Conglomerado Cafiada Honda: Prozzi 1990; G. Ramos et al. 1996), whilst carbonates are entirely lacking. Widely distributed centimetre to tens-of-metres thick, and up to 1 km long, layers of (meta-) rhyolitic to dacitic rocks are interpreted as of volcanic origin (Brodtkorb et al. 1984; Fernandez et al. 1991), whilst dykes occur west of the La Escalerilla granite (e.g., von Gosen & Prozzi 1996; von Gosen in press). The dykes were injected prior to compressive deformation and regional metamorphism of the country rocks. The age of the clastic succession is not constrained by fossils, but can be broadly compared with that of the Late PrecambrianEarly Cambrian Puncoviscana Formation s.l. of NW Argentina (e.g., Prozzi 1990; von Gosen & Prozzi 1996). In two NNE-SSW-trending strips, clastic metasediments of the Micaschist Group accompany the Phyllite Group in the southwestern part of the sierra (Fig. 1). To the west, an additional layer extends along the eastern margin of the La Escalerilla granite and further to the NNE. The Micaschist Group consists of a monotonous succession of alternating biotite-muscovite schists and quartzites with intercalated pegmatites and a few plagioclase amphibolites. In view of continuous structural and metamorphic gradation

(see below), acid meta-igneous layers, and the monotonous pelitic to psammitic composition, both eastern strips of micaschist are considered equivalent to the Phyllite Group clastic rocks (yon Gosen 1998). To the east, the micaschists continuously grade into injected muscovite-biotite schists, with quartzites and several generations of pegmatite (in part the 'Conlara Metamorphic Complex' of Sims et al. 1997). These rocks are widely distributed in the eastern part of the sierra and sometimes show a transition to migmatites and gneisses. A Late PrecambrianEarly Cambrian age is also assumed for simply deformed phyllites and micaschists with quartzites which occur between Quines and Santa Rosa. Due to poor outcrop, however, the contacts with the injected schists to the east and south, and within the monotonous succession, are unclear (simplified and generalized on the map of Fig. 1). Within the sierra, an undifferentiated basement complex can be separated into two units which are referred to here as Western and Eastern basement complexes (WBC and EBC, respectively) (see also Ortiz Sufirez et al. 1992). In the western part of the mountain range (Fig. 1), the basement rocks ('Nogoli Metamorphic Complex' of Sims et al. 1997) are injected muscovite-biotite and biotite muscovite schists or gneisses, with intercalated biotite schists and quartzites. They contain layers and lenses of amphibolite (e.g., south of San Francisco del Monte de Oro: see Merodio et al. 1978). Some of the meta-clastic rocks display different stages of migmatization. To the east of the La Escalerilla granite and adjacent micaschists, migmatitic sillimanitegarnet-biotite schists to gneisses, biotite-plagioclase gneisses, and quartzite layers comprise the EBC (in part the 'Pringles Metamorphic Complex' of Sims et al. 1997). Many areas comprise injected schists, gneisses, and quartzites with densely distributed granitic, pegmatitic, and aplitic veins and dykes on different scales. These rock types alternate with and show transitions to migmatites. Within this basement unit, lenses of mafic to ultramafic rocks and complexes with metre to kilometre scale are aligned in a NNE-SSWtrending strip (see e.g., Gonzfilez Bonorino 1961, and the 'Las Aguilas Group' of Sims et al. 1997). The most prominent occurrences are at Las Aguilas in the south and Virorco to the NNE. The mafic and ultramafic rocks include pyroxenites, norites, peridotites, hornblendites and gabbros, and are partly associated with granulites (e.g., Gonzfilez Bonorino 1961; Cucchi 1964;

EVOLUTION

OF THE SIERRA DE SAN LUIS

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Fig. 1. Simplified geological sketch map of the Sierra de San Luis based on Yrigoyen (1981), Ortiz Sufirez et al. (1992), Brogioni et al. (1994), Llambias et al. (this volume), and our own investigations. Thick vertically and horizontally ruled areas represent the Western and Eastern Basement Complex, respectively. Encircled numbers refer to different tonalite-granodiorite occurrences mentioned in the text which are: 1, Bemberg; 2, Las Verbenas; 3, Gasparillo; 4, Tamboreo; 5, Realito. Frames indicate the areas of Figs 3 and 7, whereas lines show those of profiles A-A' and B B ~ in Fig. 4.

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Sabalfia et al. 1981; Brogioni 1992; Malvicini & Brogioni 1993; Sims et al. 1997). Further north-northeastwards, the occurrences of La Bolsa and La Melada-La Gruta mostly consist of meta-gabbros with minor hornblendites and pyroxenites (e.g., Brogioni & Ribot 1994). The above units of the Sierra de San Luis are intruded by different granitoid plutons which are partly described below when their relationships with the structures is essential (see also Llambias et al. this volume).

Pre-Famatinian cycle After a penetrative foliation development (Sxl), defined by aligned biotite + muscovite, the WBC

(Fig. 1) was injected by millimetre to several metres-thick quartz-plagioclase veins and granitic sheets. This was accompanied by the growth of < l c m long muscovite plates within the country rocks. The second deformation (Dx2) led to tight Fx2 folding of the Sxl foliation and quartz-plagioclase veins on a cm to several metres scale (Fig. 2). A penetrative Sx2 foliation is the dominant planar fabric in the basement rocks. Within strips of variable width it separates Fx2 structures parallel to the axial planes. It partly records a steeply inclined to subvertical L• lineation defined by aligned biotite 4- muscovite. First magmatic layers are boudinaged and/or foliated. During Dx2, mm- to cm-thick quartz-plagioclase and aplitic to pegmatitic

Fig. 2. Simplified, schematic, and composite block diagram of fabric elements and different intrusions in the basement complex northwest of La Carolina (west of Pampa de Las Invernadas). Undeformed granite, pegmatite, and aplite dykes cut across pre-Famatinian fabric elements as well as granite injections and intrusives in biotite-muscovite to muscovite-biotite schists to gneisses. The dykes are younger than undefonned early Famatinian granites (scale is approximate).

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Fig. 3. Simplified geological map of the area northwest of La Carolina, based on interpreted aerial photographs and local field mapping (for location, see Fig. 1). The indicated sinistral and local dextral sense of shear along main fault (mylonite) zones is related to an oblique displacement.

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veins were injected parallel to, or at low angle to, the Sx2 foliation. Parallel veins were boudinaged within the Sx2 fabric. Layers and lenses of amphibolite lie within, and are affected by, the Sx2 foliation fabric, and those NW of La Carolina were interpreted as having been mafic intrusions prior to the main phase of deformation and metamorphism (Merodio et al. 1978). After the Dx2 deformation, various parts of the basement were densely injected by granitic and pegmatitic magma, leading to the development of migmatitic textures. The random growth of muscovite plates near the injections can be related to heat and fluid transfer in the country rocks. The migmatization interferes with the final stage of compressive deformation which is indicated by a local, open Fx3 folding on a metre to several metres-scale around steeply inclined to vertical axes. It is important to note that the subvertical Sx2 foliation follows a variable c. WNW-ESE to NNE-SSW trend of strike (Fig. 3) which is related to the F,,3 folding. Thus, the Bx2 axes vary in strike and record variable plunge to eastern and western directions. High-grade metamorphic overprint is indicated by the growth of biotite, muscovite, +garnet, +cordierite, and sillimanite. The latter grew pre- to syn-Dx2 and mostly was enclosed as fibrolite by muscovite plates. Migmatization during Dx3 can be related to an increase in temperature through the stages of compression, which is supported by cordierite-bearing leucosomes cross-cutting the gneissosity (Sims et al. 1997). The occurrence of staurolite and kyanite NW of La Carolina (Kilmurray & Dalla Salda 1977) cannot be directly related to this metamorphism due to the lack of detailed studies. The pre-Famatinian age of the basement complex (Llambias et al. 1996) is discussed below.

The Famatinian cycle Eastern Basement

Complex

In the basement east of the La Escalerilla granite (Fig. 1), fabrics comparable to the Dxl to Dx3 deformations of the WBC were also found with different stages of magmatic injections and migmatization. To assume a Proterozoic age for the mafic to ultramafic rocks would be in accordance to the opinions of, e.g., Kilmurray & Dalla Salda (1977), Ramos (1988), and Gonzfilez Bonorino & Llambias (1996). Recent U-Pb dating of zircons from a felsic segregation in the ultramafic rocks at Las Aguilas and of zircon rims from a spatially associated felsic orthogneiss, however, gave 478 i 6 Ma and 484 • 7 Ma,

respectively (Sims et al. 1997), which the latter authors relate to Early Ordovician intrusive activity, metamorphism, and synchronous deformation. However, this does not exclude the possibility that pre-Famatinian structural and metamorphic events exist in this complex. In contrast to the WBC, the steeply inclined to subvertical SF2 foliation has a NNE-SSW trend. Tight FF2 fold structures on a cm- to m-scale have axes with variable NNE or SSW plunge. Furthermore, wide areas record migmatitic fabrics which grade into injected sillimanitegarnet-biotite schists without muscovite. The penetrative SF2 foliation development led to boudinage of the widely distributed mafic to ultramafic lenses and mafic/ultramafic complexes. Prior to and after the compressive DF2 event, smaller magmatic bodies were injected by quartz-plagioclase and granitic veins which indicate the onset and continuation of migmatization, respectively. In some parts, an older foliation in the amphibotites, recorded by aligned long axes of amphibole, is folded around FF2 fold structures, which can be directly related to the tight FF2 folding in the metasediments. After DF2 deformation, intense migmatization in parts of the basement also affected those country rocks that contain boudins of mafic and ultramafic rocks. Isolated boudins can be found within nebulitic migmatites, recording restite textures sillimanite-garnet-biotite schists which partly display relics of tight FF2 fold structures. The formation of these migmatites interferes with FF3 fold structures around steeply inclined to vertical axes. The Ol=2 to OF3 history was accompanied by high-grade metamorphism. Granulite facies is reported from the Las Aguilas and Virorco areas (e.g., Gonz/dez Bonorino 1961; Sims et al. 1997). In the migmatitic rocks, the metamorphism is indicated by the growth of biotite, poikiloblastic garnet (also in magmatic injections), aligned and randomly oriented sillimanite, plagioclase, and K-feldspar in the absence of muscovite, whilst cordierite is described from the Las Aguilas area (e.g., Sabalfia et al. 1981; Malvicini & Brogioni 1993; Brogioni 1994; Sims et al. 1997). The growth of cm-long muscovite plates is related to later granitic and pegmatitic injections. Plastic deformation of cumulates and intercumulate material, together with partial recrystallization and mylonite fabrics (Brogioni 1992), polygonal (recrystallization) fabrics of labradoritic plagioclase (Brogioni & Ribot 1994), and the replacement of pyroxene by Mg-hornblende (Brogioni 1994) can be related to the high-temperature event within the mafic/ ultramafic complexes.

EVOLUTION OF THE SIERRA DE SAN LUIS Early Intrusions Prior to compressive deformation, the Phyllite Group and equivalent parts of the Micaschist Group were intruded by the tonalites to granodiorites of Gasparillo, Bemberg, Las Verbenas, Tamboreo (e.g., Zardini 1966; Sato & Llambias 1994; S/tnchez et al. 1996; Sato et al. 1996; yon Gosen 1998, in press; Llambias et al. this volume), and Realito (Brogioni et al. 1994), as well as additional unnamed intrusions and dykes in the western part of the sierra (Fig. 1). Intrusion resulted in contact metamorphism of the undeformed sedimentary country rocks within more or less pronounced aureoles (e.g., S/mchez et al. 1996; Sato et al. 1996; yon Gosen 1998, in press). The plutons are aligned in a long, c. N N E SSW-trending belt which, according to Sato et al. (1996) and Llambias et al. (this volume), is part of the Famatinian magmatic arc. Rb-Sr dating of the Bemberg Tonalite has given an approximate Late Cambrian to Early Ordoviclan age (c. 513 Ma errochron: Sato et al. 1996; 512 + 16 Ma: Llambias et al. this volume). U-Pb dating of zircons from the Tamboreo and Bemberg Tonalites at 470 • 5 and 468 + 6 Ma, respectively (Sims et al. 1997), suggests emplacement during the Early Ordovician and represents initial Famatinian intrusive activity. In the La Florida area, one rhyolitic-dacitic dyke cuts across the southwestern margin of the Tamboreo tonalite, whereas such cross-cutting relationships have not been found for the dykes west of the La Escalerilla granite (von Gosen in press).

Uplift and age o f the Western Basement Complex In the western part of the sierra, Famatinian tonalites and later granites (e.g., Rio Claro) intrude the WBC with sharp contacts and without direct relation to the migmatites. Muscovite granites on a scale of several metres to hundreds-of-metres sometimes display a biotite-rich marginal facies. Some enclose xenoliths of basement rocks with centimetre to metre sizes, indicating high-level stoping. NE-SW- to N-S-trending, centimetre to tens of metres thick granite and pegmatite dykes cut across the steeply inclined basement rocks together with tonalite and granite intrusions. They are crosscut by younger, c. N-S or E-W-trending pegmatite and aplite dykes. On a regional scale, the dykes record a consistent strike (Fig. 3). This shows that different orientations of the S,,2 foliation within the gneisses are related to the

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pre-Famatinian Fx3 folding event, which is confirmed by pegmatite dykes cutting across Fx3 folds on a metre scale. In the entire Phyllite Group such dykes only occur just to the east of the Gasparillo tonalite and south of Pampa de Las Invernadas. They are related to the pegmatite dykes within the tonalites and WBC. The Gasparillo tonalite intruded both the WBC and undeformed sediments of the Phyllite Group in the east. It encloses xenoliths of the basement rocks (Llambias et al. 1996) and at the northern margin cuts a large-scale Fx3 antiform. Furthermore, parts of the basement west of the La Escalerilla granite (Pancanta area), as well as adjacent parts of the Phyllite Group to the north, were intruded by tonalites and muscovite+biotite granites at the same level, without evidence for displacement along an intervening mylonite zone. All this suggests that the WBC was displaced against the consolidated deposits of the Phyllite Group prior to the onset of Early Ordovician intrusive activity. Hence, it is probable that the monotonous Late Precambrian-Early Cambrian pile of clastic sediments in the sierra covered this consolidated and steeply inclined to subvertical basement complex, which was uplifted to a high crustal level along deep-seated fault zones during pre-Early Ordovician times. From this a Proterozoic age of the basement is assumed here, although Pampean orogenesis (c. MidCambrian) cannot be excluded.

D e f o r m a t i o n and m e t a m o r p h i s m o f the Micaschist Group and equivalents East of the La Escalerilla Granite. A strip of micaschists (Fig. 1) records at least one compressive deformational event prior to the first deformation of the Phyllite Group and La Escalerilla Granite (see below). The deformation is indicated by a penetrative foliation fabric, defined by aligned biotite and muscovite, which is parallel to or cuts across (relicts of) bedding. After cessation of compressive deformation, widely distributed granitic, pegmatitic, and aplitic veins, dykes, sills, and sheets of variable extent were injected into the metaclastic rocks. The metamorphic overprint led to the growth of biotite and garnet+(fibrolitic) sillimanite. Muscovite plates and aggregates enclose fibrolite and were deformed during the later foliation development. As in the micaschists to the east, sericite aggregates probably represent andalusite converted during later compressive deformation and metamorphism.

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West and east of the Phyllite Group. The above metamorphic overprint can be broadly compared with that in the strip of micaschists of La Florida and Paso del Rey-Rio de La Carpa. There, a widely distributed pre- to partly syndeformational growth of biotite, muscovite, muscovite plates, garnet and partly (?) andalusite could be detected in the clastic successions. Fibrolitic sillimanite (not found at La Florida and north of Cerros Lagos) is mostly enclosed in muscovite plates or aggregates, as well as in up to I cm long porphyroblasts (?andalusite) that were entirely converted into sericite aggregates. Diffusional mass transfer during metamorphism is indicated by the occurrence of diffusion haloes in the micaschists of the Paso del Rey area. Planar inclusion trails (quartz grains, opaque needles and grains) within biotite and garnet + sericite aggregates :t: muscovite record different orientations with respect to SF~ planes in the Micaschist Group, and a few curved and Sshaped trails in garnet were only found within fault zones at the boundaries to the Phyllite Group. In the latter, planar inclusion trails within biotite and garnet indicate the predeformational onset of metamorphism. Thus, the present evidence suggests a variable temperature overprint which took place in the consolidated sediments prior to and partly during compressive Dw deformation. The Dw deformation led to the formation of a (partly) penetrative Svj foliation with a c. W N W ESE-trending stretching lineation (LFx) defined by sheared, flattened and elongated sericite aggregates and muscovite plates, as well as aligned biotite and muscovite. The foliation is partly related to tight FF1 fold structures and inhomogeneously affected the Paso del Rey granite and pegmatites as well as the Tamboreo tonalite. It appears to be concentrated along pegmatite and pluton margins, but also cuts across them. Interior parts of the intrusions are less deformed, or the foliation is replaced by ductile shear zones. To the east of the strip of phyllites, Svl planes are related to open, WNW-vergent FF~ fold structures within the micaschists. The folds, on a metre to tens of metre scale, become tight eastwards, and the penetrative SF1 foliation also affects pre- to (?) syn-deformational pegmatite dykes in the injected muscovite-biotite schists. On SF~ planes, the Lvl lineation gently plunges towards the ESE and is defined by aligned biotite, muscovite, < l . 5 c m long muscovite plates, and elongated muscovite-sericite aggregates. The DF~ event was followed by the injection of centimetre to tens of metres thick pegmatite dykes which led to the growth of

muscovite plates at the contacts with the country rocks. In the injected micaschists, these pegmatites are widely distributed and also some post-Dvl granite dykes occur in the east. Within parts of the eastern micaschists and injected schists to the east, the DF2 deformation is documented by open-to-tight fold structures on a centimetre to tens of metres scale. The BF2 axes trend c. E-W to NW-SE. On Sv2 crenulation cleavage planes, a c. ESE-plunging Lv2 lineation is defined by aligned biotite, muscovite, muscovite plates, and elongated sericite aggregates. In the injected schists, the gentle plunge of Lv2 can change to c. NE, and BF2 axes also strike e. NNE-SSW. During the DF2 event, older pegmatites and granite dykes were deformed. Thin layers of syn-kinematic pegmatite dykes were injected parallel to, and boudinaged between, SF2 planes. The cooling of the Rio de La Carpa granite took place during this deformation and was accompanied by the injection of up to several metres-thick pegmatite dykes in the country rocks and injected schists. The variable metamorphic overprint of the clastic suc-cession started prior to the Dvl deformation. Local growth of sillimanite within former (?andalusite) porphyroblasts suggest an increase in temperature which reached amphibolite facies conditions during the DF2 event. The climax of metamorphism is indicated by quartz-plagioclase leucosomes parallel to SF2 planes and diffuse migmatitic parts in the injected schists.

Northern part o[" the Sierra. In the Cerros Largos area and to the north, the compressive deformation in the Micaschist Group and parts of the Phyllite Group is documented by complicated fold interference patterns on a metre- to kilometre-scale, indicating at least three stages of a continuous folding process around steeply inclined to vertical axes (von Gosen 1998). In some parts, the large-scale patterns can be interpreted as sheath fold geometry which could have formed due to more intense compression in this northern part of the Micaschist and Phyllite Group, and possibly were also influenced by the shapes of (?syn-tectonic) granitoid plutons in the subsurface. The structures suggest that the compressive event affected both the phyllites and micaschists, and a clear separation of post-Famatinian fabrics was not possible. Within the phyllites to the south, however, the first deformation can be placed in the time interval following intrusion of the La Escalerilla Granite (see below). North of Cerros Largos, the successions of the Phyllite Group record continuous and

EVOLUTION OF THE SIERRA DE SAN LUIS short-distance metamorphic transitions into the Micaschist Group to the east and west which records a pre- to partly syn-Dvl growth of (?) andalusite. This also indicates that these units are lateral (and vertical) equivalents. To the east, the strip of micaschists records a gradual transition into injected micaschists. These metamorphic transitions can be explained by a variable heat transfer from underlying plutons. In the phyllites and micaschists between Quines and Santa Rosa, the first deformation is also assumed to be Famatinian in age. It is indicated by a widely distributed Svl cleavage within the steeply inclined strata. On SF~ planes, the LF1 lineation steeply plunges c. E or c. W and is recorded by aligned sericite, muscovite, and biotite. Cleavage-bedding relationships and some minor folds on a metre-scale suggest the existence of large c.W- and c.E-vergent Fv~ fold structures around c. NNE-SSW axes. In the phyllites south of Quines, however, tight FF1 folds with WNW-ESE-striking axes occur. The metamorphic overprint is indicated by the growth of sericite, muscovite, biotite, and preto syn-Dvl porphyroblasts which were partly chiastolite, but have been mostly replaced by sericite/muscovite aggregates. In the micaschists east of San Martin, the first deformation was followed by the growth of andalusite porphyroblasts that were subsequently affected by a second foliation (Ortiz Sufirez 1988). Observations within the different areas of the micaschists, injected schists, and partly also the phyllites suggest a variable, temperature-dominated metamorphic overprint with local growth of (fibrolitic) sillimanite (micaschists) and/or (?) andalusite. It started prior to the DF~ event and continued through the DF2 deformation. Eastern bounda O, of the Eastern Basement Complex. The eastern boundary of the EBC against the Micaschist Group is widely masked by thin granite and pegmatite intrusions (not depicted on the map of Fig. 1). A lamella of the Paso del Rey Granite intruded the contact between the multiply-deformed migmatites in the northwest and undeformed micaschists in the southeast. This suggests uplift of the migmatites with respect to the micaschists prior to granite intrusion. It is supported by (deformed) granites (e.g., Cruz de Carla) and pegmatite dykes within the basement complex which cut across Dvl 3 fabrics and migmatite textures, suggesting a high level of intrusion. For the Paso del Rey granite, a Mid-Late Ordovician age of emplacement was based on Rb Sr dating (454 4-21 Ma: Llambias et al. 1991). Since the granite was affected by inhomogeneous Sw shearing and low-grade

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metamorphism, which are equivalents of the events within the adjacent micaschists, the possibility cannot be excluded, however, that this date indicates the first Famatinian structural and metamorphic overprint in this part of the Micaschist Group. Hence, during and/or after this stage, some parts at least of the Micaschist Group could have been deformed in a single event combined with a slight metamorphism. North of Cerros Largos, a granite intruded the micaschist strip to the west of the phyllites. Fabrics in shear zones at the eastern contact suggest that it was emplaced during the second deformation of the country rocks. South of Paso del Rey, however, intrusive activity took place after the DF1 event within the micaschists, e.g., the La Florida granite.

Post-Famatinian magmatism, compression and metamorphism La Escalerilla Granite The intrusion of the long lamella of the La Escalerilla Granite (Fig. 1) indicates a turn in the entire tectonism. The granite intruded the WBC, Las Verbenas tonalite (Sato 1993), and undeformed clastic sediments of the Phyllite Group, and presumably followed an old fault zone. Along its western margin, the intrusion led to a contact metamorphism in the sediments (von Gosen in press). A Rb-Sr errorchron of 414Ma for the granite has been reported by Llambias et al. (this volume). U-Pb dating of zircons from the granite, however, gave an age of 4 0 3 + 6 M a , indicating a Early Devonian crystallization of the pluton (Sims et al. 1997). Based on this age, those structural and metamorphic events which affected the granite and have comparable equivalents in the surrounding units of the sierra should be younger and are described below. The structures in the Phyllite and Micaschist Group, however, are not everywhere clearly related to this late overprint (e.g., the northern part of the sierra, see above), and also in the eastern strip of phyllites a partial Famatinian overprint cannot be excluded. Structure and metamorphism Phyllite Group. The dominant structures in the Sierra de San Luis follow a NNE-SSW trend. The Phyllite Group and its contact metamorphosed equivalents were affected by a compressive D1 deformation which led to the formation of tight F I fold structures around NNE-SSW trending axes. According to way-up criteria, they

244

W. V O N G O S E N & C. P R O Z Z I

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EVOLUTION OF THE SIERRA DE SAN LUIS affected an upright sedimentary sequence. The F1 structures generally do not record a single vergence and both WNW-vergent folds and SE- to ESE-directed vergences occur (Fig. 4). The folds, with a maximum wavelength of at least several hundred metres, are associated with a penetrative, axial-plane S1 cleavage. The deformed Conglomerado Cafiada Honda lies on the eastern limb of an F 1 anticline. S 1 planes penetratively cut across rhyolitic dacitic dykes which are also folded around F1 structures. In the tonalites and in the Rio Quinto and La Escalerilla granites inhomogeneous S1 shearing was mostly concentrated along the margins, being replaced by shear zones in the interior parts. In the Phyllite Group and deformed plutons, the L1 stretching lineation strikes c.WNW-ESE. WNW-ESE compression continued locally with formation of c. W-vergent F2 crenulation folds, combined with an $2 crenulation cleavage. The WNW-ESE-striking L2 lineation is defined by aligned sericite + biotite. Western Basement Complex. Along c. N N E SSW- to N-S-trending faults, the WBC is displaced against the Phyllite Group in the east, and both contain tonalite intrusions (Figs 1 and 3). To the NW of La Carolina (Pampa de Las Invernadas), the S1 cleavage in the Phyllite Group and intercalated tonalite to granodiorite intrusive rocks is related to the F 1 folding event. It records a continuous transition into the foliation of two mylonite zones (Smy) and could be traced across these into the western basement (Fig. 5). Kinematic indicators in the mylonites demonstrate oblique sinistral uplift of the western basement block parallel to the pronounced mylonite lineation (Lmy). This is in accordance with the sense of displacement deduced for the fault zone (Fig. 3), as well as the sense of shear within some C1 shear zones in the basement (Fig. 6). A local dextral sense of shear could be detected only for parts of a NNE-SSW-trending mylonite strip at the eastern margin of the Rio Claro granite, and in one part of the main fault zone which curves around the eastern tonalite (Fig. 3). The fact that the Gasparillo tonalite, and a muscovite granite to the southwest, are not cross-cut by a mylonite zone suggests reactivation that concentrated only along those parts of the former block boundary against the Phyllite Group not sealed by intrusions. The sinistral faults partly cut through the basement (and the southwestern margin of the Gasparillo tonalite; Llambias et al. 1996) or die out to the south in the folded Phyllite Group (south of La Carolina). Except for undeformed areas with preserved pre-Famatinian fabrics, metre to hundreds of

245

metres wide strips in the WBC are inhomogeneously overprinted by the S1 foliation, which also affected the Famatinian intrusions. Foliation development started within the densely distributed pegmatite dykes. In the gneisses, the foliation can be replaced by distinct, c. N N E SSW trending, sinistral C1 shear zones (Fig. 6). An increase in shearing led to the formation of a penetrative S 1 foliation which overprinted the pre-Famatinian fabrics. La Escalerilla Granite and Micaschists to the east. The already foliated (SF1) strip of micaschists east of the La Escalerilla Granite was affected by W-vergent F1 folding and S1 cleavage formation during D1 compression in the granite and Phyllite Group to the west. The eastern margin of the La Escalerilla granite lamella is marked by a mylonite zone (Figs 1 and 4, profile B-B'). Along the mylonite, the eastern micaschists were carried northwestwards over the granite. Shear sense indicators suggest an oblique sinistral reverse faulting parallel to the mylonite lineation (von Gosen & Prozzi 1996). Within the granite, an inhomogeneously distributed S1 foliation and mostly NNW-SSE trending sinistral strike-slip faults and shear zones are related to the first deformational event and the formation of the mylonite zone at the eastern granite margin. The ductile strike-slip faults and shear zones also cut across the Las Verbenas tonalite and contact metamorphosed parts of the Phyllite Group in the west. There, they are combined with F1 folding and S1 cleavage formation. The entire deformation is related to a sinistral transpressive regime which probably was also influenced by the initial, elongate shape and orientation of the granite pluton (von Gosen in press). It led to the bending of the granite lamella together with adjacent units to the west and east (Fig. 1). Eastern Basement Complex. To the east of the La Escalerilla Granite, a steeply eastwardsinclined and thin mylonite zone displaces the gneisses and migmatites of the EBC, also northwestwards, over the strip of micaschists in the west (Figs 1 and 4, profile B-B'). There and further to the east, mylonite formation is related to the D1 event within the Micaschist Group, which led to the formation of the mylonite zone at the eastern margin of the La Escalerilla granite and compressive D1 deformation in the Phyllite Group (von Gosen in press). In the eastern part of the micaschists, compression locally continues with the formation of F2 fold structures and a ductile, sinistral C2 shear zone.

246

W. VON GOSEN & C. PROZZI

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EVOLUTION OF THE SIERRA DE SAN LUIS

247

Fig. 6. Sinistral shear zone in migmatitic gneiss at Rio Turbio, north of road from La Carolina to San Francisco del Monte de Oro (plan view). Shearing is concentrated within distinct ductile shear zones. They are replaced by penetrative S 1 foliation planes in Famatinian pegmatite dykes which cut across the pre-Famatinian S~2 foliation.

The EBC is displaced along a N N E - S S W trending belt of mylonites (Fig. 1), referred to as the 'Faja milonitica La Arenilla' by Ortiz Sufirez et al. (1992). In the southwestern part of the sierra, it consists of one thick mylonite (e.g., in the mafic/ultramafic La Melada complex; Brogioni & Ribot 1994). In the northern and southern continuation, the mylonite splits into several branches which are accompanied by additional, parallel and thinner mylonites (Fig. 7). Regionally, the mylonite foliation (Stay) strikes N N E - S S W . In the outcrops, it mostly steeply dips to the east or is vertical; a gradual change to a steep westerly dip only occurs locally. Smy follows the Famatinian SF2 foliation (Fig. 8). Since no F1 folding could be detected outside the mylonite zones, this also shows that slip during displacement of the different basement blocks mostly used pre-existing Sv2 foliation planes. In the southern part of the Virorco complex, bending of the older foliation is related to a large-scale fold structure (Gonz~tlez Bonorino 1961), interpreted here to be of early Famatinian age and cross-cut by mylonites.

Several thin mylonite zones partly surround the Virorco occurrence and cut through the Las Aguilas complex (Fig. 7; compare also, e.g., Gonzfilez Bonorino 1961; Cucchi 1964; Sabalfia et al. 1981; Brogioni 1992). In the strip of basement rocks to the west of the thick mylonite zone, a more intense S1 foliation occurs parallel to Sv2 planes. The basement to the east is inhomogeneously affected by Sv2 parallel shearing or a more intense S1 foliation within distinct strips between and towards mylonite zones. The mylonites displace basement blocks with enclosed boudinaged layers of mafic and ultramafic rocks which were injected by granitic and pegmatitic melts before (Fig. 8; see also above). They also cut across the migmatites, the latest-stage fabrics of the Dr3 deformation, and contain clasts of the migmatites. Kinematic indicators (a- and ~5-clasts, shear bands, S/C fabrics) show uplift of the eastern basement blocks (Figs 7 & 8). In many cases, this was combined with a variable sinistral component parallel to the mylonite lineation (Lmy), which plunges c. SE to c.S. Lmy is defined by aligned sillimanite, biotite, • elongated quartz

248

W. V O N G O S E N & C. P R O Z Z I

Fig. 7. Simplified map of mylonite zones within the Eastern Basement Complex and at the eastern margin of the La Escalerilla granite, based on field mapping and interpreted aerial photographs (for location see Fig. 1). Mafic and mafic/ultramafic complexes (grey) are simplified and only the larger occurrences are indicated on the map. Lower hemisphere, equal area stereoplots indicate the maximum of mylonite foliation planes (Stay; great circle) and mylonite lineations (Lmy, circle) as well as fold axes related to the mylonite formation. Numbers of measurements are indicated in brackets. Depicted relative sense of shear (black and white half circles) is based on shear sense indicators (cr- and &clasts, shear bands, S/C fabrics). G, La Escalerilla Granite; M, Micaschist Group.

EVOLUTION OF THE SIERRA DE SAN LUIS

249

Fig. 8. Mylonite strips in migmatitic basement gneisses and schists at Riocito (north of La Melada, east of asphalt road to La Carolina). Quartz-plagioclase injections are folded and/or boudinaged within the pre-Famatinian Sx2 foliation. They cut across amphibolite layers and during boudinage they were sheared and flattened at the margins. The mylonite foliation (amy) follows the Sx~ foliation fabric or cuts across under a small angle (left). Shear sense indicators (&clasts, shear bands, SIC fabrics) indicate an oblique sinistral uplift of the eastern blocks parallel to the mylonite lineation (Lmy). (Schematic block profile.)

and quartz-feldspar aggregates, and tails of quartz on feldspar clasts. The oblique sinistral component of displacement is comparable with that found at the eastern margin of the La Escalerilla Granite, at the eastern boundary of the WBC, and is supported by sinistral strike-slip shearing in the La Escalerilla Granite and to the west (see above). From E to W, however, the sinistral component becomes more important. This shows that it is not a local phenomenon but was of regional significance during compression (see below). Within the mylonites, elongated quartz ribbons surround quartz and feldspar clasts (and aggregates) which are bent, kinked, and mechanically twinned (plagioclase). Ductile deformation was accompanied by widely distributed recrystallization of quartz. Recrystallization of feldspar took place at clast boundaries and/or along shear planes. Old garnet poikiloclasts are affected by the foliation whilst smaller grains or rims grew during and/or after shearing. Sillimanite is partly bent, kinked, disrupted, and converted into sericite. With respect to the fabrics, the mylonites can be compared with

the mylonite at the eastern margin of the La Escalerilla Granite, where coarser feldspar recrystallization grains also occur. This also accounts for the static growth of quartz and feldspar recrystallization grains after cessation of shearing, whilst adjacent layers indicate continuing ductile deformation and dynamic recrystallization. These observations suggest that mylonite formation, which started under at least higher greenschist facies conditions during uplift, continued at lower temperatures. It is possible, however, that mylonitic shearing partly reactivated some older (Famatinian) mylonite strips, since Sims et al. (1997) report high-temperature mylonites from, e.g., the Las Aguilas area. The greenschist facies fabrics suggest that mylonite formation took place coeval with the shearing at the eastern margin of the La Escalerilla granite. It is thus directly connected with the compression which led to F1 folding and S1 cleavage formation within the Phyllite Group and equivalent fabrics of the Micaschist Group. Hence, uplift along the intra-basement mylonite zones east of the La Escalerilla Granite can be interpreted as

250

W. VON GOSEN & C. PROZZI

having accommodated compression which led to folding and cleavage formation within the 'cover sediments' (e.g., Phyllite Group). Micaschists east and west of the Phyllite Group. Between the EBC and the Micaschist Group to the east, a c. NE-SW-striking, sinistral mylonite zone was found NW of Paso del Rey. The mylonite partly affects the Paso del Rey granite and pegmatites, but to the NE it cuts through the migmatites. As at the eastern margin of the WBC, it can be assumed that displacements only affected those parts of former block boundaries which were not sealed by granitoid and pegmatite intrusions. Along high-angle reverse faults, the deformed Micaschist Group was carried over the central strip of the Phyllite Group (e.g. Fig. 4, profile A-A'). As in the case of the above mylonite zones, they are related to the overall c. W N W ESE compression in the sierra, but they record W- and E-directed reverse senses of displacement, in accordance with different fold vergences. In the Micaschist Group of the Paso del Rey and La Florida areas, an SE- to ESEvergent F 1 folding with an S 1 crenulation cleavage becomes more intense towards the faulted contact with the Phyllite Group in the east. At the eastern faulted contact of the phyllites, a steeply eastwards inclined S1 foliation occurs within the micaschists of Rio de La Carpa. In the south, this contact is marked by a c. 500 m wide zone of into intensely foliated (S1) micaschists, where Svl planes are overprinted (e.g., Rio de La Cafiada Honda). According to c.W-vergent F1 structures, S/C fabrics, and shear bands, the micaschists are carried westwards over the phyllites (the 'Rio Guzm~n Shear Zone' of Sims et al. 1997). Only at Rio Quinto could an oblique sinistral component be found. A NNE-SSW-trending shear zone, cutting across the Phyllite Group north of Cerros Largos, indicates uplift of the eastern block. However, it is related to the first and second deformations with complex folding and is presumably of Famatinian age (see above). In all areas, the along-strike continuations of these fault lines are unclear, and the faults are omitted on the map of Fig. 1. Injected micaschists and phyllites to micaschists between Quines and Santa Rosa. Only within a few outcrops, Fv2 folds within the injected micaschists are overprinted by open to tight, c. WNW-vergent F I folds on a decimetre to metre scale, which do not record an S 1 cleavage. After folding, the sequence was injected by up to

several metres thick pegmatite dykes. These are partly related to local migmatization due to postkinematic granitoid intrusion in the eastern part of the sierra. Migmatization also affected those rocks that record Fv2 fold structures. In the phyllites to micaschists between Quines and Santa Rosa only a few c. ESE- or c.WNWvergent, open F1 crenulation folds on a centimetre to several metres scale were found; they are combined with an S 1 crenulation cleavage. Regional metamorphism and age. In the metaclastic sediments, the regional metamorphic overprint, accompanying the compressive deformation, is indicated by recrystallization of biotite, quartz, 4-feldspar, muscovite, and muscovite plates, as well as the growth of biotite, chlorite, sericite/muscovite, +garnet. Former (? andalusite) porphyroblasts are converted into sericite/muscovite +chlorite aggregates. Partial replacement of garnet by biotite/chlorite, and of biotite by chlorite, is also attributed to the final static stage of metamorphism. Comparable greenschist facies fabrics from granitic and tonalitic intrusions are described by, e.g., Sato (1993), Brogioni et al. (1994), and von Gosen (1998, in press). In the basement complexes, new muscovite, sericite and biotite 4-garnet grew. Older garnet and sillimanite may be partly to entirely replaced by biotite and sericite, respectively, whereas feldspar contains sericite. From the mafic to ultramafic complex of La Melada, the effects of the metamorphic overprint are described by, e.g., Brogioni & Ribot (1994). Based on ductile deformation and dynamic recrystallization of feldspar, a relatively high grade of greenschist facies can be estimated for the regional metamorphism ( T > c. 450~ The overprint, however, was modified by the formation of late migmatites in the eastern part of the sierra, which are related to post-kinematic intrusions. Relict staurolite, indicating amphibolite facies conditions, could be detected only at two localities of the micaschists in the southwestern part of the sierra. Whilst staurolite in the La Florida area can be related to the intrusion of the La Florida granite, south of La Carolina it grew statically after the second deformation within the micaschists. Hence, the precise areal distribution of amphibolite facies overprint is unclear. Locally occurring kyanite in isolated blocks of the micaschists east of the La Escalerilla granite, large staurolites entirely replaced by sericite aggregates east of Cerros Largos, and staurolite NW of Cerros Largos (Kilmurray & Dalla Salda 1977) could also be Famatinian in age.

EVOLUTION OF THE SIERRA DE SAN LUIS In many parts of the different units, regional greenschist facies metamorphism statically outlasted compression. However, deformation locally continued during the final stage of heating. This is exemplified also in the mylonite zones where ductile deformation and dynamic recrystallization of quartz 4- feldspar continued, whereas cessation of displacements along distinct parts of the same mylonite, or within another mylonite, led to static growth of recrystallized quartz and feldspar (blastomylonites). K-Ar data from biotites of the Paso del Rey and Rio de La Carpa granites gave ages of 372 + 20, 381 -t- 13, and 391 4- 9 Ma (Varela et al. 1994), older than a single date from the La Escalerilla granite (353+15Ma: Linares & Latorre 1973). Since these granites were affected by compressive deformation and greenschist facies metamorphism, the data can be interpreted to reflect the closure of the K-Ar system in Devonian time. This is supported by Ar/Ar data from sericite in the mylonitic fabric of the 'Rio Guzman Shear Zone', which indicates an age between 360 and 350Ma (Sims et al. 1997). Together with the above indications, this would imply a probable Devonian timing of the post-Famatinian compressive deformation and regional metamorphism in the sierra, which is in accordance to the interpretation of Sims et al. (1997).

251

during Mid-Cambrian times (e.g., Acefiolaza et al. 1978; Miller 1984; Willner et al. 1987; Fig. 9) as the result of collision with the Arequipa-Antofalla craton (Ramos 1988). This contrasts with the situation in the Sierra de San Luis where no indications for Pampean compression and metamorphism have yet been reported from the Phyllite Group and equivalents. Due to the lack of radiometric data indicating a Pampean event, the WBC is interpreted as Proterozoic in age, consistent with its uplift prior to intrusion of early Famatinian tonalites and granites. Hence, it is reasonable to assume that the sediments of the Phyllite Group and equivalents equally covered an older basement complex in a passive margin setting. The deposition of the Conglomerado Cafiada Honda can be related to submarine debris flow from a fault escarpment during turbiditic sedimentation. Recent data from Rapela et al. (this volume) show that the Pampean cycle of the Eastern Sierras Pampeanas (Sierras de C6rdoba) was caused by an Early to Mid-Cambrian collision between continental terranes. This supports the existence of a 'Pampean Terrane' (Ramos 1988), which collided with the Rio de La Plata craton in the east during the Pampean cycle (Baldo et al. 1996), and contrasts with the model of Kraemer et al. (1995), who proposed a collision in Proterozoic times prior to deposition of the Puncoviscana Formation.

Geotectonic implications Famatinian and post-Famatinian Pre-Famatinian

cycle

For the low grade meta-clastic sediments of the Phyllite Group in the Sierra de San Luis a minimum initial thickness of c. 3.5 km can be estimated (yon Gosen & Prozzi 1996). These rocks can be compared with single, isolated occurrences of low grade rocks in the provinces of C6rdoba, Catamarca, and La Rioja (compare, e.g., Lucero Michaut 1979; Prozzi & Ortiz Sufirez 1994; G. Ramos et al. 1996). Allowing for gradation into higher grade metamorphic units, this suggests a thick pile of clastic sediments and a widespread depositional area outside the present Sierra de San Luis. This could indicate that it was an equivalent of, or connected to, the Puncoviscana trough of NW Argentina, if one accepts a comparable depositional age in both realms. The clastic successions of the Puncoviscana basin were deposited on the western passive margin of Gondwana during the Late Precambrian-Early Cambrian interval (e.g., Je~ek et al. 1985) and were affected by the Pampean Orogeny

cycles

During the Famatinian cycle, the formation of a magmatic arc at the western margin of the Pampean Terrane is indicated by the rise of tonalite, granodiorite, and granite plutons in the sierra (Sato et al. 1996; Llambias et al. this volume). The initial tonalites intrude the undeformed sediments of the Phyllite Group and parts of the Micaschist Group, as well as the WBC which had already been uplifted. The magmatic arc probably continues northwards in the Sierras de Los Llanos, Malanz~in and Chepes (Fig. 9; see also V. A. Ramos et al. 1996, Llambias et al. this volume) where intense subduction-related plutonism occurred in the earliest Ordovician-Caradoc interval (Pankhurst et al. 1996, this volume). The mafic and ultramafic rocks in the EBC of the Sierra de San Luis intruded the metasediments during the early Ordovician, synchronously with regional deformation, and represent a significant heatsource for the high-temperature conditions (see Sims et al. 1997). This suggests that these events were slightly older than, or occurred during,

252

W. VON G O S E N & C. P R O Z Z I

Fig. 9. Simplified geological map of the NW Argentine Andes with main geotectonic units based on published maps and data (Willner et al. 1987; Miller & Neugebauer 1990; Ramos 1991; Simon & Rossello 1990; L6pez & Toselli 1993; Neugebauer & Miller 1993; Vujovich 1994; Lopez et al. 1996a, b; V. A. Ramos et al. 1996; Schetselaar et al. 1996; Toselli et al. 1996) and our own investigations. Boundaries between geotectonic units are inferred; my, mylonite zone; arrows schematically indicate tectonic transports; encircled numbers refer to locations mentioned in the text: 1, Sierra de E1 Gigante; 2, Sierra de Guayaguas; 3, Sierra de La Huerta; 4, Sierra de Valle F6rtil; 5, Sierra de Las Minas; 6, Sierra de Chepes; 7, Sierra de Los Llanos.

EVOLUTION OF THE SIERRA DE SAN LUIS the onset of arc plutonism. The latter can be related to continuous, east-directed subduction at the western margin of the Pampean terrane. According to the plate tectonic situation, this was combined with eastward drift of the Precordillera (Cuyania) terrane towards the Gondwana margin. The timing of compressive deformation and variable metamorphism in the Micaschist Group and its equivalents suggests onset in the (?) MidLate Ordovician interval. This also applies to parts of the Phyllite Group and injected schists in the eastern part of the sierra, although early Ordovician events cannot be excluded. It is important to note that most parts of the Phyllite Group were not affected by compression, suggesting a variable areal distribution of compressive effects and metamorphism during the Famatinian cycle. A few observations show that uplift of the EBC took place prior to and after the first deformation within different areas of the micaschists, and was presumably combined with the onset of intra-basement mylonite formation. After intrusion of the La Escalerilla Granite, WNW-ESE compression under greenschist facies conditions led to folding of the cover sediments and probably was accommodated by the formation of mylonite zones in the EBC, leading to uplift of different blocks. Oblique reverse faulting with a sinistral component in the western part of the sierra, a strike-slip reactivation of former block boundaries (e.g., WBC), and displacements of the Phyllite Group over the Micaschist Group (Fig. 10) are all related to this compressive event, which probably is Devonian in age. With respect to the main geotectonic units and suture zones, Fig. 9 shows a preliminary interpretation. Based on the extent of the Precordillera (Cuyania) terrane (cf. V.A. Ramos et al. 1996), a main suture can be inferred between the Sierra de Valle F+rtil-La Huerta and the Sierra de E1 Gigante west of the Sierra de San Luis (southern continuation of the Valle F6rtil lineam e n t ) - see Schmidt (report during the 1995 Penrose Conference), Astini et al. (1995), Schmidt et al. (1995) and V. A. Ramos et al. (1996). Vujovich (1994) reports mylonites and NW-directed transport in the Sierra de La Huerta, whilst Simon & Rossello (1990) describe a minor sinistral component from mylonites in the Sierra de Guayaguas. The abundance of marble in the Sierra de E1 Gigante, which does not occur in the Sierra de San Luis proper, and also a different deformational and metamorphic history with a general c. E-W trend of the main structures (see Gardini 1993), suggests that the former represents an isolated block that can be

253

interpreted as part of the Western Sierras Pampeanas (Gardini et al. 1996). Schetselaar et al. (1996) mention counter-clockwise bending of the main structures in the western margin of the Sierra de E1 Gigante (cf. Schmidt et al. 1995), related to kink folds and ductile shear in the outcrops, and interpret these features as the expression of a sinistral Early Palaeozoic strikeslip fault. The fault was later followed by an Andean thrust (Schmidt et al. 1995). Several localities with mylonite have been described from the eastern margin of the Famatina System (e.g., P6rez et al. 1991; L6pez & Toselli 1993; L6pez et al. 1996a, b; Miller & Neugebauer 1990; Neugebauer & Miller 1993; and Toselli et al. 1996). Studies of kinematic indicators (e.g., P6rez et al. 1991; L6pez et al. 1996a, b; Toselli et al. 1996) have shown that mostly W- to NW-directed, reverse displacements of the eastern blocks took place (simplified in Fig. 9). A Devonian timing for the formation or reactivation of some of these mylonite zones was described by L6pez et al. (1996a) and Toselli et al. (1996) although some are older (e.g., P~rez et al. 1991). Furthermore, Kntiver & Miller (1981) and Willner (1983) have reported a Silurian age of deformation and metamorphism from the Sierra de Ancasti. In the map of Fig. 9, the western margin of the Famatina System is inferred. If one accepts that the Valle F6rtil lineament and its southern continuation represent the (relic of the) main terrane boundary, then a thin slice of Pampean Basement is intercalated between the Precordillera (Cuyania) terrane and the Famatina System, and probably extends to the south (with respect to fault lines; compare also interpretations of, e.g., Acefiolaza & Toselli 1988 and Baldis et al. 1989). The Famatina System would represent a sliver of magmatic arc pinched within the Sierras Pampeanas basement (Miller & Neugebauer 1990). Taken together, these observations, as well as those from the Sierra de San Luis, show that oblique sinistral reverse to partly sinistral strike-slip displacements also occurred at the inferred contact between the Precordillera (Cuyania) terrane and the Sierras Pampeanas basement complex (Valle F6rtil lineament). This suggests that general E-W to WNW-ESE compression between the different geotectonic units was not necessarily accommodated only by pure W-directed reverse displacements along mylonite zones. With respect to the curved outer shape of the eastern boundary of the Precordillera (Cuyania) terrane, the sinistral components are presumably the effect of the colliding terrane indenting the Sierras Pampeahas basement to the east. The crucial point is the

254

W. VON GOSEN & C. PROZZI

r,n

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~ ~.~

EVOLUTION OF THE SIERRA DE SAN LUIS timing of the docking of the terrane with the Pampean terrane in the east. Our studies in the Sierra de San Luis make it possible to assume that post-Famatinian compression, affecting all units, was related to collision. Conclusions

In the Sierra de San Luis, the different stages of the structural and metamorphic evolution can be separated as follows and are schematically summarized in Fig. 10. (1) The pre-Famatinian cycle is indicated by at least three phases of compressive deformation within the Western Basement Complex (Dxl-3), accompanied by a high-grade metamorphism leading to the formation of migmatites. These events are interpreted as of Proterozoic age. (2) The timing of deposition of a thick clastic sequence (later Phyllite Group and equivalents) is compared with that of the Late PrecambrianEarly Cambrian Puncoviscana Formation of N W Argentina. The sediments probably covered the consolidated and steeply inclined Western Basement Complex and were intruded by rhyolitic-dacitic dykes. (3) After uplift of the Western Basement Complex, intrusion of tonalites in the undeforreed clastic sequence are related to the Famatinian magmatic arc and assigned to the Early Ordovician. They also support the assumed timing for the evolution of the basement as well as the deposition of the clastic sequence. This igneous activity is related to an E-dipping subduction zone at the western margin of the Pampean terrane, probably linked to the E-directed drift of the Precordillera (Cuyania) terrane. The early Ordovician intrusion of mafic and ultramafic melts in the Eastern Basement Complex were combined with deformation and high-grade metamorphism (Sims et al. 1997). (4) To a variable extent, Famatinian deformation(s) occurred within the clastic cover sequences and affected parts of the Phyllite Group. The age of the variable metamorphic overprint, which partly reached amphibolite facies conditions, can only be assumed to fall in the (?) Mid-Late Ordovician interval. (5) Post-Famatinian compression led to folding and cleavage formation within the clastic cover sediments, inhomogeneous foliation development within the Famatinian plutons, and displacement of different basement blocks along mylonite zones, which also occur at reactivated older block boundaries. Kinematic indicators from mylonite zones and from shear zones in the La Escalerilla Granite suggest a sinistral c o r n -

255

portent of displacement during c . W N W - E S E compression in the western part of the sierra. Uplift of basement blocks along mylonite zones was probably combined with folding in the cover sediments and partial displacement of the Micaschist Group over the Phyllite Group. (6) Compression was accompanied, and mostly outlasted, by greenschist facies metamorphic conditions. Based on the age of the La Escalerilla Granite (Sims et al. 1997) and a few additional radiometric data, this phase of deformation and metamorphism probably took place in the Devonian. Post-Famatinian deformation and metamorphism are assumed to have been related to the colliding Precordillera (Cuyania) terrane in the west. An oblique sinistral component during compression, found in the western parts of the sierra and supported also by observations from outside the sierra, can be interpreted as the effect of the shape of the indenting terrane. We are very grateful to E. J. Llambias and A. M. Sato (La Plata) for many discussions about the granitoids, their relations to country rocks, and the evolution of the Sierra de San Luis on fieldtrips and elsewhere. Discussions with M. de Brodtkorb (Buenos Aires), W. Buggisch (Erlangen), C. Costa, C. Gardini, A. Ortiz Su/trez, G. Ramos (San Luis), S. Delpino, L. V. Dimieri (Bahia Blanca), P. Gonzhlez, C. Rapela (La Plata), and V. A. Ramos (Buenos Aires) helped to understand the evolution of the sierra and surroundings. We are grateful also to those colleagues who helped and supported us in solving logistic problems. Thanks to R. J. Pankhurst, J. Sims and an anonymous reviewer for critical comments and textural suggestions. Over several years, our work in the Sierra de San Luis was enabled by grants from the CONICET, Universidad Nacional de San Luis, German Academic Exchange Service (DAAD), and German Research Foundation (DFG, Proj. Bu 312/29-1 and 29-2). Cars were provided by the Universidad Nacional de San Luis, Universidad Nacional de La Plata, and the Proyecto de Ciencia y T&nica U.N.S.L. ('Mapa GeolSgico de la Sierra de San Luis y areas circunvecinas'). These supports from all above institutions are gratefully acknowledged. References

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U-Pb, Th-Pb and Ar-Ar geochronology from the southern Sierras Pampeanas, Argentina: implications for the Palaeozoic tectonic evolution of the western Gondwana margin J. P. S I M S 1, T. R. I R E L A N D 2'3, A. C A M A C H O 1, P. L Y O N S 1, P. E. P I E T E R S 1, R. G. S K I R R O W 1, P. G. S T U A R T - S M I T H

1 & R. M I R O 4

1Australian Geological Survey Organisation, PO Box 378, Canberra A C T 2601, Australia 2 Research School of Earth Sciences, Australian National University, Canberra ACT, Australia 3 Dept of Geological & Environmental Sciences, Stanford University, CA 94305, USA 4 Servicio Geol6gico Minero Argentino, Av. Poeta Lugones, Nueva C6rdoba, Argentina Abstract: New SHRIMP zircon and monazite 2~

and 2~ ages on structurally controlled units and 4~ step-heating ages from shear fabrics, define three distinct regional tectonic events in the southern Sierras Pampeanas. The first, the Pampean orogeny, involved closure of a late Neoproterozoic basin on the western margin of Gondwana. New rims on detrital zircons and concurrent monazite growth suggest that the metamorphic peak was attained by c. 530 Ma. The second event, the Famatinian orogeny, marks the initiation of eastward-dipping subduction on the western Gondwana margin, and may represent a continuation of the earlier Pampean event. Metasedimentary rocks from the Sierras de San Luis have zircons with a predominantly Early Cambrian detrital age, indicating a Pampean source. The metamorphic peak in these rocks was contemporaneous with the emplacement of felsic, mafic and ultramafic rocks at c. 480 Ma in a collisional setting. Monazite ages and limited new zircon growth in the metasedimentary rocks suggest that the Famatinian orogeny had ceased by about 450 Ma. This correlates well with a 450460 Ma Ar-Ar age for late shearing in the southern sierras of La Rioja province. The third tectonic event, the Achalian orogeny, involved W-directed compression and emplacement of multiple, voluminous, granite intrusions. Deformation during this event was partitioned between discrete shear-zones and regions of open to tight folding. The shear zones alternate between W-directed thrusts and NNW-trending, sinistral shear-zones. Ar-Ar data from the low-grade shear fabrics indicate that transpressional deformation continued through most of the Devonian. The western part of the palaeocontinent of Gondwana comprised numerous cratonic blocks that amalgamated during various events from the Proterozoic through to the earliest Palaeozoic (e.g., Unrug 1992). Critical to our understanding of the make-up of the supercontinent and times of accretion and dispersal, are constraints on the ages of formation and deformation of the various cratonic blocks. In this respect, the southern Sierras Pampeanas terrane in N W Argentina is somewhat enigmatic. Placed between the Grenville-aged (c. 10001200 Ma) basement of the Precordillera in the west (e.g., Kay et al. 1996) and the Archaean to Palaeoproterozoic Rio de La Plata craton in the east (e.g., Cordani & Brito Neves 1982), there are a number of conflicting views about the age of: (1) sedimentation; (2) initial deformation and metamorphism in the basement rocks; and, (3) the duration and timing of subsequent events. In most plate reconstructions, the high-grade basement rocks of the Sierras Pampeanas are

considered to be Neoproterozoic in age with significant deformation dating from the Brasiliano orogeny (640-570Ma; e.g., Bond et al. 1984; Ramos 1988; Brito Neves & Cordani 1991; Unrug 1992). Ramos (1988) suggested that this deformation in the Sierras Pampeanas involved westward-directed subduction of a marginal basin at the western edge of the Rio de La Plata craton, beneath the Pampean terrane. Alternatively, Acefiolaza & Miller (1982) used stratigraphical evidence to suggest that the metasedimentary rocks of the Sierras Pampeanas were deposited in a marginal basin, the Puncoviscana Trough of Je2ek et al. (1985), adjacent to the Rio de La Plata craton during the latest Proterozoic to earliest Cambrian. Acefiolaza & Miller (1982) further suggested that the terrane then underwent orogenesis during a Mid-Late Cambrian event that they termed the Pampean Orogeny (after Acefiolaza & Toselli 1976), followed by a period of sedimentation during the Late Cambrian.

SIMS,J. P., IRELAND,T. R., CAMACHO,A., LYONS,P., PIETERS,P. E., SKIRROW,R. G., STUART-SMITH,P. G. & MIRO, R. 1998. U-Pb, Th-Pb and Ar-Ar geochronology from the southern Sierras Pampeanas, Argentina: implications for the Palaeozoic tectonic evolution of the western Gondwana margin. In: PANKHURST,R. J. & RAPELA,C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 259-281.

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SIERRAS PAMPEANAS: PALAEOZOIC EVENTS Regardless of whether the initial deformation occurred during the Neoproterozoic (i.e., the Brasiliano orogeny) or Cambrian (i.e., the Pampean orogeny), most authors agree that subsequent events may be grouped into a prolonged event termed the Famatinian orogeny (after Acefiolaza & Toselli 1976). Dalla Salda et al. (1992) suggested that this orogeny involved a Cambrian, pre-collisional subduction-phase, which culminated in the collision of Laurentia and Gondwana in the middle Ordovician (c. 480-460 Ma). This was followed by granitic plutonism in the Siluro-Devonian. This threefold separation of orogenic events is, in part, consistent with the interpretation of Rapela et al. (1990) who suggested that granites within the Sierras Pampeanas could be subdivided on the basis of composition into three groups (G l-G3) related to the tectonic history. The first group (G1) was thought to be related to highgrade metamorphism during the Neoproterozoic to Cambrian, thereby indicating a link with either the Brasiliano orogeny of Ramos (1988) or the Pampean orogeny of Acefiolaza & Miller (1982). The second group (G2: syn-D2) of Rapela et al. (1990) is related to the collision of Laurentia with Gondwana, which, in the scenario of Dalla Salda et al. (1992), produced an orogenic belt with high-grade metamorphism and widespread granite intrusion. The third intrusive group (G3) was related to a further subduction episode that developed well to the west of the Pampean terrane in the late Palaeozoic (Rapela et al. 1990). Previous geochronology in the southern Sierras Pampeanas has been predominantly concerned with the ages of intrusive units, and there are few reliable ages for the associated metamorphic rocks (e.g., Linares & Gonz/tlez 1990). Knfiver & Miller (1981), however, have reported early Palaeozoic Rb-Sr ages for metamorphic basement in the Sierras de Ancasti in the central Sierras Pampeanas. They suggested that metamorphic events in that area could be identified in early Cambrian and middle Ordovician times, with a static recrystallization event occurring in the Silurian (Kn~ver & Miller 1981). An Early Cambrian Rb-Sr age for a metamorphic event in the southern Sierras Pampeanas has been confirmed by Rapela & Pankhurst (1996), and Rapela et al. (this volume). This paper presents new geochronological data, relevant to the tectonic history of the southern Sierras Pampeanas, obtained during a two year, integrated geoscientific mapping

261

programme, undertaken by the Australian Geological Survey Organisation (AGSO) in collaboration with the Servicio Geoldgico Minero Argentino (SEGEMAR). This mapping program was undertaken in three separate areas (Fig. 1) in the southern Sierras Pampeanas, covering an area of about 27 000 km 2, and was supported by new geophysical (aeromagnetic and radiometric) data sets. The three areas were chosen to provide a semi-continuous cross-section through the major tectono-stratigraphical packages comprising the terrane. We report on: (1) the provenance and age of the basement metasedimentary rocks through examination of detrital zircon populations; (2) the metamorphic history of the terrane determined by dating new zircon rims on detrital cores and concurrent monazite growth; (3) the crystallization age of important ultramafic rocks and felsic orthogneiss through dating of magmatic zircons; and, (4) the age of late shearing episodes based on Ar-Ar isotopic analyses of micas in structural fabrics. These data constrain the duration of the Pampean and Famatinian orogenies to the Early Cambrian and EarlyMid-Ordovician (Tremadoc-Caradoc) respectively. Furthermore, the data define an important new deformational event in the southern Sierras Pampeanas, the Achalian orogeny, that spanned the Devonian Period.

Geological setting Located in NW Argentina, the southern Sierras Pampeanas are a series of ranges which comprise crystalline basement intruded by diverse Palaeozoic igneous rocks (Fig. 1). These rocks were exposed during Miocene to Recent compressional (Andean) tectonics (e.g., Jordan & Allmendinger 1986). In general, the high-grade basement rocks consist of pelitic to psammitic units, with a number of subordinate carbonatedominated lithologies; metamorphic facies ranges from greenschist to marginal granulite. Structurally overlying the basement rocks are several low-grade, metasedimentary belts, whose metamorphism post-dates that in the highergrade basement (e.g., Sims et al. 1997). These low-grade belts include the San Luis Formation in the Sierras de San Luis (e.g., Ortiz Sufirez et al. 1992), the Mermela Group (Baldo et al. 1996) and the younger Tasa Cuna Formation (Gonz/tlez & Acefiolaza 1972) in the Sierras de Cdrdoba,

Fig. 1. Generalized geology of the southern Sierras Pampeanas. Dashed boxes show the areas covered by the AGSO-SEGEMAR project. Locations of Figs 2, 3 and 4 are indicated by the solid boxes.

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J. P. SIMS E T A L .

and the MalanzS.n Formation in the southern sierras of La Rioja Province (Furque 1968). Even though these low-grade rocks were deposited after the high-grade metamorphism of the basement, both the San Luis Formation and the Mermela Group record episodes of compressional deformation that resulted in folds, shear zones and development of a slaty cleavage (Baldo et al. 1996; von Gosen & Prozzi 1996). The intrusive rocks of the southern Sierras Pampeanas fall into three dominant age groups: (1)

(2)

(3)

Early Cambrian I- and S-type intrusives that are largely restricted to the northern Sierras de Cdrdoba and the Sierra Norte de C6rdoba (e.g., Rapela & Pankhurst 1996; Lira et al. 1997; Lyons et al. 1997; Rapela et al. this volume); Early Ordovician intrusives occurring dominantly in the southern sierras of La Rioja Province and in the western Sierras de San Luis (e.g., Pankhurst et al. 1996; Pieters et al. 1997; Sims et al. 1997; Pankhurst et al. this volume); Voluminous Devonian intrusives exposed mainly in the Sierras de Cdrdoba and the eastern Sierras de San Luis (e.g., Ortiz Sufirez et al. 1992; Demange et al. 1996; Sims et al. 1997) and which are also interpreted as occurring under shallow Cenozoic cover adjacent to the southern sierras of La Rioja Province (Pieters et al. 1997).

Additionally, mafic and ultramafic rocks occur in both the Sierras de C6rdoba and Sierras de San Luis. The significance of these rocks is that they have previously been interpreted as indicative of Neoproterozoic subduction-related processes. For example, Ramos (1988) has interpreted the mafic and ultramafic rocks in the Sierras de San Luis as having formed in a back-arc extensional setting during the Neoproterozoic Brasiliano orogeny. Similarly, Escayola et al. (1996) have interpreted two belts of ultramafic rocks in the Sierras de C6rdoba as ophiolites from both back-arc and mid-ocean ridge settings.

Sample descriptions In order to constrain the timing of deposition and age of subsequent orogenic events that affected the metamorphic basement of the southern Sierras Pampeanas, a range of rock-samples have been selected for analysis. Four samples of metasedimentary rocks from the high-grade basement ( P P l l l a , PL063, PL147 and JS129c) were selected for detrital zircon analyses.

Monazite from two of these samples (PL063 and JS129c) was also analysed and their (metamorphic?) ages compared with the ages of zircon overgrowths found in the corresponding detrital zircon populations. Two igneous samples representing spatially associated ultramafic and felsic intrusives were selected (JSOSOf and JSO79e) to constrain emplacement ages for those bodies. Additionally, since the ultramafic rocks are inferred to be the heat source of metamorphism in the surrounding metasediments (Sims et al. 1997), these samples should constrain the local age of the metamorphic peak. Five samples from mylonite zones (ACOllb, JSO86f, JS13Ob, PP192b and PL230b) were also selected to determine the timing of mica growth in shear fabrics that overprint the high-grade basement. The location of the samples is shown in Figs 2, 3 and 4, and they are further described below.

Metasedimentary

rock samples

P P l l l a is from a cordierite schist of the Olta metamorphic complex (Pieters et al. 1997) within the Sierra de Chepes (Fig. 2), in La Rioja Province. The Olta metamorphic complex forms the basement of the southern sierras of La Rioja Province and consists largely of psammitic and pelitic metasedimentary rocks ranging from greenschist-facies phyllites and schists to migmatites. These metasedimentary rocks are intruded by granites of the Chepes igneous complex, which have yielded U-Pb zircon crystallization ages in the range 495-470Ma (Pieters et al. 1997; Pankhurst et al. this volume; Fig. 2) and Rb-Sr ages in the range 470450Ma (Pankhurst et al. 1996). PL063 is from a weakly layered, migmatitic, biotite-

rich gneiss of the Pichanas metamorphic complex (Lyons et al. 1997) in the northern Sierras de C6rdoba (Fig. 3). This complex consists predominantly of cordierite- or garnet-bearing paragneiss and migmatite, with minor marble and amphibolite units. A number of highly peraluminous granite bodies with early Cambrian Rb-Sr (520 :t: 5 Ma; Rapela & Pankhurst 1996) and U-Pb (c. 527 Ma; Lyons et al. 1997) isotopic ages intrude the Pichanas metamorphic complex (Fig. 3). PL147 is from a well foliated, migmatitic, garnetbearing gneiss of the Quilpo formation (Lyons et al. 1997), also in the northern Sierras de C6rdoba, though lying approximately 65 km to the east of PL063 (Fig. 3). The Quilpo formation is one of several carbonate-rich metasedimentary units within the southern Sierras Pampeanas and consists predominantly of interlayered pelitic and psammitic gneiss, marble, calc-silicates and amphibolite. JS129c is from a layered, biotite-garnet-sillimanite gneiss of the Pringles metamorphic complex (Sims et al. 1997) in the Sierras de San Luis (Fig. 4). The main rock-types in the Pringles metamorphic complex are pelitic and psammitic gneiss and schist, which are intruded by mafic and ultramafic bodies and lenses. This complex transects the Sierras de San Luis (Figs 1 & 4) and has a unique geophysical signature that

SIERRAS P A M P E A N A S : P A L A E O Z O I C EVENTS

263

Fig. 2. Generalized Palaeozoic geology of the Sierras de Chepes, La Rioja Province (after Pieters et al. 1997), showing location of geochronology sites (with sample numbers) discussed in the text. Heavy line-work represents various generations of ductile and brittle faults. Additional age data are U-Pb zircon ages for various intrusive phases of the Chepes Igneous Complex from Pieters et al. (1997). N-S and E-W grid has 10km spacing.

distinguishes it from similar rock-units of high metamorphic-grade and complicated structural histories to both the east and west (Sims et al. 1997). This complex is also structurally interleaved with low-grade metasediments of the San Luis Formation, which is intruded by intermediate granitoids with c. 470 Ma ages (Sims et al. 1997; Fig. 4).

Igneous rock samples Zircons from two igneous rocks that intrude the Pringles metamorphic complex have been analysed in order to determine their crystallization ages. Sample JSO80f is a plagioclase-rich, orthopyroxene-bearing, felsic segregation within a pyroxenite phase of the Las

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Fig. 3. Generalized Palaeozoic geology of the northern Sierras de Cdrdoba, Cdrdoba Province (after Lyons et al. 1997), showing location of geochronology sites (with sample numbers) discussed in the text. Heavy line-work represents various generations of ductile and brittle faults. Light line-work within geological units represents boundaries of compositional variation. The additional age is a U-Pb zircon age for the El Pildn Granite from Lyons et al. (1997). N-S and E-W grid has 10kin spacing. Aguilas intrusion (Gonzalez Bonorino 1961). This intrusion is one of a number of mafic and ultramafic bodies (the Las Aguilas group of Sims et al. 1997) that occur within the Sierras de San Luis (Fig. 4) and were interpreted by Ramos (1988) as emplaced during the Neoproterozoic Brasiliano orogeny. JSO79e is a representative sample of several small, recrystallized, felsic intrusive bodies that are spatially, closely associated with the Las Aguilas group. This sample consists predominantly of recrystallized quartz and feldspar, poikiloblastic garnet and biotite which wraps around the garnet and forms a late anastomosing, retrograde fabric.

Shear-zone samples Samples ACOl lb, JSO86f, JS13Ob, PP192b and PL230b are from low-grade shear-zones that truncate the various metamorphic and igneous complexes. In the Sierras de San Luis, there is a series of these shear zones that have variable shear-sense and have equilibrated at different crustal levels. Sample ACOIlb is from a pegmatitic vein within a narrow (metre-scale) N-trending shear zone, in the central-west of the Sierras de San Luis (Fig. 4). Kinematic indicators within this shear-zone show an east-side up shearsense, on a steeply plunging mineral lineation. The

mineral assemblage of kyanite-staurolite-muscoviteisillimanite within the pegmatitic vein suggests that this shear-zone formed at pressures of about 4 kbar or greater. Sample JSO86fis from a mylonitized, NNW-trending pegmatite (c. 5 m wide) in the southwest of the Sierras de San Luis (Fig. 4), with a subhorizontal stretching lineation and sinistral shear-sense indicators. Orientated, fine-grained muscovite and euhedral garnet, have overgrown the original pegmatite assemblage during mylonitization. Sample JS130b is from the Rio Guzman shear-zone (Sims et al. 1997), a major NNE-trending, greenschist-facies shear-zone that transects the Sierras de San Luis. This shear-zone, which is prominent on Landsat-TM, aeromagnetic and radiometric imagery, is up to 3 km wide and separates high-grade basement to the east from low-grade metasedimentary rocks of the San Luis Formation (Fig. 4). The mylonitic fabric, which dips steeply to the east, is defined by muscovite-chlorite-quartz+euhedral magnetite, with a near-vertically plunging elongation lineation defined by quartz ribbons and slickenlines. Shear-sense indicators show east-side up. Sample PP192b is from a mylonitized granite on the margin of a N-trending shear-zone that forms the contact between metasedimentary rocks of the Olta metamorphic complex and intrusives of the Chepes igneous complex in the Sierra de Chepes (Fig. 2). The shear fabric is defined by an anastomosing muscovitebiotite foliation, with recrystallized quartz forming a

SIERRAS PAMPEANAS: PALAEOZOIC EVENTS

265

Fig. 4. Generalized geology of the Sierras de San Luis, San Luis Province (after Sims et al. 1997), showing location of geochronology sites (with sample numbers) discussed in the text. Light line-work within geological units represent boundaries of compositional variation. Additional age data are U Pb zircon ages from Sims et al. (1997); 468 + 6 is for the Bemberg Tonalite, 470 :k 5 Ma is for the Tamboreo Granodiorite and 403 + 6Ma is for the Escalerilla Granite. INS and E W grid has 10 km spacing.

266

J. P. SIMS E T A L .

weak mineral lineation. The mylonitic fabric dips steeply to the east with east-side up shear-sense. Sample PL230b is from the N to NNW-trending Guamanes Shear Zone (Martino 1993; Lyons et al. 1997) in the Sierras de C6rdoba (Fig. 3). The shear zone consists of several generations of proto- to ultramylonitic fabrics that are dispersed over a 7 km wide zone, overprinting medium- to high-grade metasedimentary rocks, including the Pichanas metamorphic complex. The mylonitic fabrics are generally moderately to steeply east-dipping, with a steeply plunging lineation locally defined by biotite and quartz ribbons. Reactivation of the shear zone is associated with chlorite hematite assemblages in zones parallel to the earlier biotite fabric. The shear-sense of both generations of mylonitic fabrics in the Guamanes Shear Zone is east-side up.

Analytical methods U-Pb

and Th-Pb

For zircons and monazites, separates were obtained from crushed, bulk rock samples using standard heavy liquid and magnetic separation techniques. The grains were then mounted with pieces of the standard zircons, SL 13 and AS3, while monazites were mounted with the Maffra monazite standard. The mounts were then polished to expose midsections. Mounts were photographed first in reflected light at low magnification to create sample location maps, and then in both transmitted and reflected light at high magnification for grain characterization. In addition, SEM-based cathodo-luminescence (CL) imaging was carried out to detect any internal structure not evident in the optical images. Ion microprobe analyses were carried out using the SHRIMP I (Sensitive High Resolution Ion Micro-Probe) at the Research School of Earth Sciences (RSES), Australian National University. The advantage of SHRIMP over conventional U-Pb zircon dating is that petrographically selected areas of complex, zoned and structured zircons (or monazite) can be analysed in situ (e.g., Williams 1992; Fitzsimons et al. 1996). Zircon data were reduced according to the methods described by Muir et al. (1996) and Ireland & Gibson (1998). In zircon, the calibrated 2~ was normalized to the mean 2~ measured for the standard (of known age and hence Pb/U), giving Pb/U ratios for the unknowns. The primary U-Pb standard zircon for this work, AS3, is derived from a gabbroic anorthosite from the Duluth Complex, Minnesota. Zircons from this rock have been shown to be consistently concordant with an age of 1099 + 1 Ma (PaGes & Miller 1989). U and Th concentrations are determined relative to those measured from the 572Ma SL13 standard which is homogeneous in terms of U and Th concentrations to within 20 per cent. Th/U ratios are calibrated relative to the measured 2~176 of the standard. AS3 standard zircon chips with concentrations around 200ppm to 400 ppm were used for Pb/U normalization. Monazite analyses were normalized to the 421Ma Maffra

monazite standard (Ireland & Gibson 1998). For monazite, an additional calibration for Pb/Th versus ThO/Th was carried out, allowing a Th-Pb age determination. The U and Th concentrations of the Maffra monazite grains are highly variable and so no concentrations for monazites are reported. Zircon ages are used in two different ways in this work. The metasedimentary rocks are analysed for detrital-zircon age spectra so as to allow a view of possible provenance contributing to the deposition centre. For these spectra, the ages of the individual zircons are pooled (by way of a probability density function) and the peaks are interpreted as likely source ages. Common Pb is removed from the analyses by monitoring the 2~176 ratio with the isotopic composition of common Pb calculated from the Cumming & Richards (1975) growth model for the inferred age of the crystal. For detrital zircons, it is important to have both 2~ and 2~176 ages of individual zircons. The 2~ ages have an accuracy of c. 2% regardless of age, whereas the precision of the 2~176 ages are a function of age and U concentration. In order to limit bias associated with a user selection of 2~ or 2~176 age, an approach based on the weighted mean of the two ages is adopted. If the MSWD of the resulting mean is less than 10, the weighted mean age is adopted. If the resulting age has a high MSWD, the two ages are examined in order to assess the discordance. In most cases, the difference in age is a result of a discordant 'old' grain and for these the 2~176 age is adopted. The remaining cases are for young grains for which an inappropriate common Pb correction has been applied; for these grains the 2~ age is adopted. While any single zircon age could be affected by Pb loss or mixed zones of a crystal, the coincidence of ages required to create the peaks in the age spectra suggests that the inferred ages are, most probably, geologically meaningful. However, the extraction of age components can be difficult when there are overlapping populations and in this case the peaks were deconvolved using the MIX programme of Sambridge & Compston (1994). For magmatic rocks of Palaeozoic age, the 2~ 2~ age is not well determined in ion probe analyses because of the low concentrations of 2~ and 2~ in the zircons and the small amount of material analysed. For example, only 3fg (i.e. 3 x 10-15 g) of 2~ is contained in the volume of a typical analysis of the SL 13 standard. Thus for Palaeozoic rocks, the 2~ is insensitive and we choose to correct for common Pb by extrapolating the measured 2~176 and 2~ to concordia. The isotopic composition of common Pb is calculated from the Cumming & Richards (1975) growth model for the inferred age of the sample. The fraction of common 2~ f2~ can be obtained from a lever rule. The use of another Pb isotopic composition would have a negligible effect on the magmatic ages since most of the uncorrected data points are close to concordia. The coherence of the population is assessed by taking the weighted mean of the individual Z~ ages. The mean-square of the weighted deviates, MSWD, should approach unity for a large population and should not exceed a critical

SIERRAS PAMPEANAS: PALAEOZOIC EVENTS value for the appropriate number of degrees of freedom for smaller populations. If the MSWD is excessive, then outliers can be assessed. The data sets are variously plotted on probabilitydensity diagrams and Tera-Wasserburg diagrams. In general, the sediments are only plotted on the probability-density diagrams. These figures are generated by assuming a unit area Gaussian for each datum and then summing contributions for all the data in 1 Ma bins. It should be noted that the Tera-Wasserburg plots are for uncorrected Pb compositions. Uncertainties for the error bars on the Tera-Wasserburg plots are at one standard deviation (lcr). Errors in ages of the peaks are reported at one standard error in the m e a n (lffm). For the final ages, the final error in the standard (which is a systematic error) is summed in quadrature with the peak age and the errors are reported at two standard error (2Crm).

Ar-Ar For muscovite and biotite, sample concentrates were prepared from crushed bulk samples by the Australian Mineral Development Laboratories (AMDEL) using heavy liquid separations and a Frantz Isodynamic Separator. Irradiation of mica separates for 4~ analysis was carried out in the H I F A R reactor of the Australian Nuclear Science and Technology Organisation. All mineral samples were irradiated for 336 hours, with the biotite standard GA 1550 (age of 97.9Ma; McDougall & Roksandic 1974) in the central position as described by McDougall & Harrison (1988). Cadmium sleeves, 0.2mm thick, were placed inside the sample cans to minimize interference from thermal neutrons. Sample cans were rotated 180~ three times

267

through the irradiation in order to minimize the effects of the neutron flux gradient in HIFAR. The micas were analysed for isotopic argon concentrations at the RSES Ar-Ar facility. The technique involved progressively increasing the temperature from 500 ~ to 1450~ with the gas from each extraction step exposed to Zr-A1 getters to remove active gases. Argon was subsequently analysed using a VG Isotech MM1200 gas source mass-spectrometer operating in static mode. Corrections for argon produced by neutron interactions with calcium and potassium were made using the factors determined by Tetley et aL (1980). The 4~ abundance and decay constants of Steiger & J/iger (1977) were used. Errors are reported at two standard deviations (2am).

Results and interpretation A s u m m a r y o f the results f r o m samples analysed o n S H R I M P is listed in Table 1, whereas A r - A r ages are s h o w n in Table 2. T h e d a t a sets o f analyses are l o d g e d as a S u p p l e m e n t a r y P u b l i c a t i o n N o . S U P 18118 ( 1 6 p p ) a n d are available f r o m the British L i b r a r y D o c u m e n t S u p p l y Centre, B o s t o n Spa, W e t h e r b y , W Y o r k s L S 2 3 7 B Q or f r o m the G e o l o g i c a l Society Library.

U - P b dating o f detrital zircons Z i r c o n s f r o m P P l l l a are typically detrital with rounded form and no apparent overgrowths. On

Table 1. Summary of U-Pb zircon and Th-Pb monazite ages Sample No

Latitude (S)

Longitude (W)

Rock type (Unit)

Mineral

Result/age (Ma)

PPllla

31.10656

66.5474

Zircon

Gondwana detrital

PL063

30.76801

65.28507

Cordierite schist (Olta metamorphic complex) Biotite-muscovite gneiss (Pichanas metamorphic complex)

PL147

30.91162

64.61434

Garnet-bearing gneiss (Quilpo formation)

JS129c

32.98101

66.07376

JS079e

33.12184

66.13525

Gondwana detrital 561 + 10 531 + 10 526 + 11 Gondwana detrital 585 -t- 16 529 + 10 (Gondwana) detrital c. 460 451 4- 10 484 4- 7

JS080f

33.12306

66.13645

Garnet-sillimanite gneiss (Pringles metamorphic complex) Felsic orthogneiss (Pringles metamorphic complex) Felsic segregation in pyroxenite (Las Aguilas group)

Zircon cores Zircon rims Zircon rims Monazite Zircon cores Zircon rims Zircon rims Zircon cores Zircon rims Monazite Zircon

Zircon

478 + 6

Errors quoted are 2orm.

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J. P. SIMS ET AL.

Table 2. Summary of Ar-Ar step heating age Sample

Latitude (S)

Longitude (W)

Rock type (unit)

Mineral

Step heating age (Ma)

Type of spectrum

JS086f

33.26173

66.17059

Muscovite

375 • 1

Flat

JS 130

33.02794

65.93205

Sheared pegmatite Mylonite

Muscovite

351 362

AC01 lb

32.88154

66.08597

Muscovite

366 • 2

Continuously rising Flat

PP192b

31.10252

66.33944

Muscovite

452-459

PL230b PL230b

31.04803 31.04803

64.82461 64.82461

Biotite Muscovite

358 • 2 365-451

Kyanite-bearing shear-zone Mylonitized granite Mylonite Mylonite

Continuously rising Flat Continuously rising

All errors quoted are 2o%.

the other hand, the zircons from samples PL063 and PL147 are generally clear and equant to elongate in outline and many of the elongate grains are structured with zoned cores and overgrowths of clear, light brown zircon. These rims have low cathodo-luminescence suggesting high U concentrations while the rounded and broken cores are typically zoned with higher luminescence (Fig. 5a, b). The zircon grains from sample JS129c are generally clear and equant in outline. A less dominant population of elongate to equant prismatic crystals with pyramidal terminations are also present. Most grains in this sample are zoned and some have structural cores and narrow overgrowths (Fig. 5c). While the detrital zircons of P P l l l a can be analysed readily for a detrital zircon age spectrum, the presence of rims on the zircons from the other three samples creates the difficulty of whether the rims are metamorphic (i.e. the rims were produced during the metamorphism of the rock) or whether the rims are related to an igneous (or metamorphic) event prior to the deposition of the sediment. In the first instance we will examine the age spectra from the 'cores' of the detrital zircons in samples PL063, PL147, and JS129c. Then we will examine in more detail the youngest detrital peak in conjunction with the zircon rims from PL063 and PL147. P P l l l a . The detrital zircon signature of 47 zircons from P P l l l a (Fig. 6a) has a dominant contribution from c. 500-700 Ma, a less distinct peak around 1000-1200Ma, and scattered ages from c. 1600 to 3400Ma with most of these ages between 1600 and 2100 Ma. This pattern is familiar in terms of Neoproterozoic to Early Palaeozoic sediments around the Gondwana margin and is found in Australia, New Zealand,

and Western Antarctica (e.g., Williams et al. 1991; Ireland et al. 1995; Gibson & Ireland 1996; Ireland & Gibson 1998). The zircon data presented here indicate a similar continental protolith for this sediment from the Sierras Pampeanas. The youngest detrital zircons from P P l l l a show a bimodal distribution with the main peaks at c. 550 and 620 Ma (Fig. 6b). PL063. Thirty four analyses were carried out on 23 zircons from PL063, comprising 14 rims and 20 cores (or whole grains). The 20 cores show a fairly typical Gondwana pattern (Fig. 6c) as exemplified by P P l l l a . The Neoproterozoic peak ranges from 600 to 700Ma; the Grenvillian peak (c. 1000Ma) is well developed, and there are two older grains at 1600 and 2500 Ma. The cores are resolved from the rims in terms of age (Fig. 6d). The youngest cores are around 620 Ma while the rims show an age range from 480 to 590Ma. Apart from the youngest analysis, which appears to have lost Pb, the rims show a bimodality in their age distribution (Fig. 6d). The two ages that are derived from peak de-convolution of the rim analyses are 531 + 10Ma (2Crm) and 561 + 10Ma (2O'm). PL147. The probability density distribution for PL147 (Fig. 6e) is very similar to that for PL063; with a dominant 600-700 Ma peak, a Grenville component and two older grains. The main peak in the cores (Fig. 6f) is spread over c. 600-650Ma; a single analysis below 550Ma is possibly the result of Pb loss. The main peak of the rims is around 530Ma with some older grains as well at c. 580Ma and one around 650 Ma (Fig. 6f). The two main peaks from the rim analyses de-convolve to give two ages of 5294- 10Ma (2Crm) and 585• 16Ma (2Crm) respectively.

SIERRAS PAMPEANAS: PALAEOZOIC EVENTS

269

Fig. 5. Examples of cathodo-luminescence images from structured zircon samples: (a) PL063, (b) PL147, (e) JS129c, (d) JSO79e. Detrital zircons (a-c) are overgrown by narrow rims (indicated by white arrows) with low luminescence. Inherited cores (in d) with high luminescence (indicated by black arrows) are overgrown by magmatic zircon with low luminescence.

JS129c. The distribution of'core' detrital ages for JS129c (Fig. 6g) shows a dominant Neoproterozoic to Early Cambrian peak at around 530 Ma, while the Grenville contribution is very subordinate and there are no older grains. Several grains that are younger (Fig. 6h) are interpreted as due to Pb loss. Several grains are older than the main peak (600-750 Ma) and these are likely due to the presence of the Gondwana signature as a minor component. Only four rims were analysed in JS129c, predominantly because of their narrow width (typically c. 10 #m) and the comparatively large size of the primary beam (20 x 30 #m) used in this work. These four (wide) rims all show ages under 500Ma. The weighted mean of these four analyses is 4 5 9 + 5 M a (l~rn) with an MSWD of 4.3. The MSWD is high, but the number of analyses is low. Rejection of either the high

value or the low value could be made but the added precision would not affect the interpretation of the results. The rims are suggestive of formation at c. 460 Ma. Under certain circumstances, the cut-off in the youngest detrital age population can be used to infer a maximum deposition age for the sediment. The geological limitations to this paradigm are: (1) that there was limited disturbance younger than the depositional age that could have caused Pb loss, and (2) that the cut-off is indeed a maximum since it requires that the zircons which were being deposited, had formed, essentially, coevally (such as tuffaceous zircons). For three of these samples, there are clearly overgrowths on zircons which may have affected their U P b systematics. Nevertheless, the main peaks in all the samples are closely related temporally. The youngest detrital grains in

270

J. P. SIMS E T AL.

= 0"201

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cores

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___________ PL063 cores n = 20

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3500

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450

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0.40

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.hA..., .... ,...,I .... , .... l ooo 1500 2o0o 25oo 30o0 3500

Age (Ma)

0.00 400

45o

500

550

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Age (Ma)

Fig. 6. Summary of the detrital zircon data from basement units in the southern Sierras Pampeanas (Sierras de C6rdoba, southern sierras of La Rioja Province and Sierras de San Luis: (a) & (b) P P l l l a from the Olta metamorphic complex, La Rioja Province (Fig. 2), (c) & (d) PL063 from the Pichanas metamorphic complex, C6rdoba Province (Fig. 3), (e) & (f) PL147 from the Quilpo Formation, C6rdoba Province (Fig. 3), (g) & (h) JS129c from the Pringles metamorphic complex, San Luis Province (Fig. 4). Data are presented as cumulative probability plots, first showing the full range of zircon ages for each sample and then, the detail in the region of the main detrital peak(s). The detailed plots also include the analyses from zircon rims, which are shown with a dashed line. Ages for the de-convolved rim peaks are shown with 2(7 errors.

SIERRAS PAMPEANAS: PALAEOZOIC EVENTS PP111a are c. 550Ma in age, the rims in both PL063 and PL147 are about 530Ma, and the main peak in JS129c is also about 530 Ma. These zircon growth periods reflect different geological events and histories. The youngest detrital grains must be eroded from their source and deposited in a sedimentary pile. The narrow range in ages from the main peaks in the four samples indicates that whatever the geological history, various geological events affecting the zircon systematics happened over a relatively short interval of time between c. 550 and 500Ma. The rims on zircons from PL063, PL147 and JS129c may reflect zircon formation in the protolith (which was then eroded and deposited), or in situ formation during metamorphism of the sedimentary rock. For unravelling the geological history, the nature of these zircon rims is probably the most important parameter. The rims typically have very low Th/U ratios compared with the 'core' analyses. This requires that there is another phase, such as monazite, removing Th from the system during the crystallization of zircon. During prograde metamorphism of pelitic compositions, monazite becomes

unstable and breaks down to REE oxide-phases, however, it re-stabilizes under upper-amphibolite conditions (Kingsbury et al. 1993). In this case, co-precipitation ofmonazite and zircon results in complementary chemistries. Alternatively, monazite and zircon may be cogenetic in magmatic systems such as S-type granites. However, in this case, zircon appears to crystallize first because the Th/U ratio of these zircon rims is in the range of other magmatic zircon (Muir et al. 1996). Therefore, detailing the age of monazites in these rocks, which have obtained sufficient metamorphic grade to recrystallize monazite, would allow us a possible diagnostic as to an in situ metamorphic growth of zircon, as opposed to inherited igneous (or metamorphic) zircons.

U - P b and T h - P b dating o f monazites PL063. Monazites from sample PL063 are generally pale yellow and rounded. U - P b and Th-Pb data are plotted in Tera Wasserburg (Fig. 7a) and cumulative probability (Fig. 7b) diagrams. While the U - P b data appear to be

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271

age: 5 2 6 _+ 11 M a (2ore)

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700

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400

450

500

550

600

650

700

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Fig. 7. Tera-Wasserburg and cumulative probability plots for monazite analyses from basement units in the southern Sierras Pampeanas (as per Fig 6): (a) & (b) PL063, (e) & (d) JS129c. Individual samples in (a) & (c) are plotted without correction for common Pb. Radiogenic 2~ compositions are calculated by projection of the data along the dashed line, which represents a mixing line to the common Pb composition of Phanerozoic crust (Cumming & Richards 1975), onto the concordia. Analyses not included in the final age calculation for the sample are shown as open circles. Cumulative probability data are presented for both Th-Pb and U-Pb ages.

J. P. SIMS ET AL.

272

bimodal (Fig. 7b), the MSWD of the peak at 1.6 is not indicative of more than one component. Also, the form of the Th-Pb peak is consistent with only one age. The Th-Pb age is 526+ 11 Ma and has a higher precision than the U - P b age mainly because of the low U in the Maffra standard and the resulting larger uncertainty in the Pb/U ratio used for normalization. The monazite age is slightly younger than the zircon age for this sample although the 2a error bounds overlap. The monazite data show that metamorphism of this sample occurred at around 530 Ma.

Comparison o f zircon and monazite data An age of c. 450 Ma for monazite growth in JS129c is consistent with the limited amount of data obtained on the zircon rims for this sample at c. 460 Ma. This suggests that the sedimentary protolith was metamorphosed through upper amphibolite facies with concomitant growth of monazite and zircon rims. This seems to be supported by a comparison of the Th/U ratios between monazite, zircon 'cores' and zircon rims for this sample (Fig. 8). Zircon 'cores' for JS129c have a mean Th/U ratio slightly less than 1.0, whereas the zircon rims, which have similar ages to the monazites, have Th/U ratios an order of magnitude lower. The monazite data for PL063 are also consistent with the age of the youngest zircon rims in the same sample, at c. 530Ma. And as for JS129c, these rims have a Th/U ratio an order of magnitude lower than the 'core' zircons (Fig. 8). This also seems to suggest that zircon and monazite formed concomitantly during a metamorphic event. It is interesting to note that the older zircon rims in this sample are easily distinguished from both the younger rims and the 'core' analyses. If the sedimentary protolith

JS129c. Monazites from sample JS129c are also generally pale yellow and rounded. U-Pb data are plotted in Fig. 7c and d. Each of the U - P b and Th-Pb data sets has three outliers. Two analyses in each are low, suggesting Pb loss, but one grain has U - P b and Th-Pb ages on the high side of the peaks. It is unclear whether this is an analytical artefact or represents some inheritance from either an earlier (metamorphic or detrital) monazite. The remaining seventeen analyses give mean U-Pb and Th-Pb ages of 452+ 12 (2O-m) and 451 + 10Ma (2crm), respectively.

'

I

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9 9

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_ _ 600

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9

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.

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-

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i

",,=.

/

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~ '

'

450

500

Age

(Ma)

550

~ 9 .

0.001 600

500

550

i

,

P P 1 1 1 a zircon Th/U l , , i , ~

600

650

700

Age (Ma)

Fig. 8. Plots of Th/U ratio versus age for samples PL063, PL147, JS129c and PPll la, showing analyses (that fall within the age-range shown) of zircon cores, zircon rims (where available) and monazites (where available). See text for further discussion.

SIERRAS PAMPEANAS: PALAEOZOIC EVENTS was metamorphosed (in situ) to yield these zircon rims, the subsequent high-grade metamorphism at 530 Ma has largely reset the corresponding monazites. In such a scenario, the nature of the resetting would probably have been through recrystallization because the temperatures required for volume diffusion of Pb through monazite in such a short time would probably result in zircon resetting or crystallization as well. Alternatively, if metamorphism occurred in the sedimentary source-rocks, erosion, deposition, and metamorphism must have occurred within c. 30Ma. With the data alone, there is little to distinguish between these two options. The age of the monazites in PL063 is also the same age as the zircon rims in PL147. Unfortunately, only PL063 and JS129c were processed for monazite and so we have no data for the age of monazite growth in PL147. The low Th/U ratios for the zircon rim analyses in comparison to the 'cores' for PL147 (Fig. 8), would suggest, though, that monazite was also forming simultaneously in this rock. The older zircon rim ages, however, are indistinguishable from the 'core' analyses in terms of their Th/U ratios and may suggest an igneous source for these rims.

273

variable and range between 127ppm and 1220ppm with Th/U ratios ranging from 0.20 to 1.37. Four U-Pb analyses are older than the main population (Fig. 9a) and have been omitted to form a satisfactory mean. Their exclusion leaves 14 analyses defining a weighted mean 2~ age of 4844-7Ma (2O-m), which is interpreted as the magmatic crystallization age for this felsic orthogneiss.

JSO8Of. The zircons from JSO80f are dominated by equant sub-round crystals some of which show pyramidal terminations. The grains are mostly clear and zoned with no apparent cores. CL imaging shows large-scale sector zoning with a finer, superimposed, oscillatory zoning and no apparent cores. Zircon U-Pb data are shown in Fig. 9b. Uranium concentrations are highly variable and range from 243 ppm to 1900ppm but there is only a narrow range of Th/U ratios from 0.18 to 0.29. All 15 U-Pb analyses are combined to give a weighted mean 2~ age of 4 7 8 + 6 M a (2Crm), which is interpreted as the magmatic crystallization age for the Las Aguilas intrusion.

Ar-Ar dating o f micas Spectra relating to 4~ step heating experiments carried out on muscovite and biotite that grew during mylonitization are given in Fig. 10. Data from the step heating analyses show that the first few steps and the last two steps generally have Ca/K ratios considerably higher than the remainder of the steps. These analyses are considered anomalous (e.g., Dickin 1995) and the derived ages are not considered to have a major geological significance.

U-Pb dating of igneous zircons JSO79e. The zircons from JSO79e are dominated by elongate and equant sub-round crystals, some of which show pyramidal terminations. The grains are mostly clear, zoned and contain abundant fluid inclusions. CL imaging shows that rims with well-developed, oscillatory zoning have developed around pre-existing, rounded cores (Fig. 5d). Uranium concentrations are

0.10

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Fig. 9. Tera-Wasserburg plots of zircon analysis from the Sierras de San Luis: (a) JSO79eand (b) JSO8Of.Plotting details as in Fig. 7.

J. P. SIMS ET AL.

274 400

I

Ia

400

400

400

390

390

390

380

380

~ 380

380

370

370

~ 370

370

360

360

~ 360

360

350 1.0 380

350

AC011b Muscovite

390

I ]

366 + 2

I

350 0.2

O. 380

0.4

0.6

0.8

JS130b Muscovite ]

t~

370

370

[ b

JS086f Muscovite]

O. 470

I

375 + 1

0.2

0.4

0.6

0.8

PP192b Muscovite

Jd

350 1 0 470

460

,..,,460 m

J

350

360

350

~450

450

~ 440

440

<

m

34O ~ 430

340

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0.2

0.4

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0.8

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420

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0.2

0.4

0.6

0.8

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I

480 PL230b Muscovite ]

] e 480

440

440

~

U__.__

358+2_ I

[f.]

r

'-'="---'-'m.y

360

g

V

400

I 360

320

320

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400 e~

O.

<

_2-

360

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320 0.

0.2

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0.8

320 1.0

Fraction 39Ar released

.

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0.6

,

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Fraction 39Ar released

Fig. 10. Ar-Ar step-heating spectra for mica samples from shear fabrics in the southern Sierras Pampeanas: (a) A COllb (muscovite), (b) JSO86f (muscovite), (c) JS130b (muscovite), (d) PP192b (muscovite), (e) PLO23b (muscovite), and (f) PL230b (biotite). Data used for calculation of plateau ages in samples ACOllb, JSO86fand PL230b(biotite) are indicated by the horizontal bar. Three of the samples, ACOllb, JSO86f and PL230b(moTITE), produced spectra which were taken to be flat overall (Fig. 10a, b, f) and produced well-defined ages. The apparent mean ages for these samples are 3 6 6 + 2 M a (2am), 375 + 1 Ma (2O'm) and 358 + 2 Ma (2Crm), respectively (Table 2). The remaining three analyses, JS13Ob, PP192b and PL230b(MuSCOVITE) produced continuously

rising age spectra without well defined age plateaux (Fig. 10c, d, e). JS130b shows an age range from c. 351 to c. 362 Ma, PP192b from c. 452 to c. 459Ma and PL230b(MuSCOWTE) from c. 365 to c. 451 Ma (Table 2). These spectra may reflect the timing of when the different domains closed to isotopic diffusion during cooling. In this case, the maximum apparent age may be interpreted as the maximum age estimate for the time of

SIERRAS PAMPEANAS: PALAEOZOIC EVENTS initial closure. And the geological evidence would suggest that this was the case for both JS130b and PP192b as the muscovites in these samples is part of the stable, shear-fabric, mineral assemblage. Alternatively, the spectra may reflect partial argon loss during a subsequent event. For example, the muscovite in sample PL230b is considered from petrographical evidence to be inherited from a previous metamorphic fabric. Thus, the rising spectra in this sample would be consistent with partial argon loss.

Discussion Integration of the new geophysical data sets with detailed mapping (Lyons et al. 1997; Pieters et al. 1997; Sims et al. 1997) and the geochronology data presented here permit a reinterpretation of the tectonic history of the southern Sierras Pampeanas. A summary of this reinterpretation is shown in Fig. 11 and is discussed below.

Depositional ages and provenance

The detrital zircon patterns we have obtained for the metasediments from the southern Sierras Pampeanas suggest an overall Gondwana margin affinity with a pattern in the main peaks not dissimilar to other occurrences from Neoproterozoic-Early Palaeozoic localities in western Antarctica, New Zealand, and eastern Australia. The three samples from Sierra de Chepes ( P P l l l a ) and Sierras de Cordoba (PL063 and PL147) have detrital patterns dominated by this type of material. On the other hand, JS129c from Sierras de San Luis has only a minor (complex) sedimentary component while the main detrital peak has probably formed from the weathering and re-deposition of a Cambrian-aged igneous protolith. The main detrital peaks from P P l l l a , PL063, and P L I 4 7 have 600-650 Ma ages, while P P l l l a has a smaller detrital peak at c. 550Ma. Both PL063 and PL147 have produced two ages for rims on the detrital 'cores'. In both of these samples, the younger rim ages of c. 530 Ma may be interpreted with some confidence as an age for in situ metamorphism. The older rim ages, however, may have three obvious interpretations. These are: (1) that these older zircon rims indicate an earlier in situ metamorphic event and hence a maximum depositional age of c. 600 Ma; (2) that the older zircon rims are detrital, hence indicating a maximum detrital age of c. 560 Ma; or, (3) that one of the zircon rims is detrital

275

while the other is due to in situ processes. We prefer the second alternative, that the rims are detrital, largely because of the structural continuity and stratigraphical similarities between the Pichanas and Cruz del Eje metamorphic complexes, and the Quilpo Formation in the Sierras de Crrdoba (Lyons et al. 1997; Fig. 3). If the third interpretation were correct, it would imply that a significant age boundary existed between the two samples, PL063 and PL147. Though a major structural boundary does exist, in the form of the Guamanes Shear Zone, our Ar-Ar data indicate this to be a younger (late Devonian) structural feature (Table 2). Furthermore, if the first alternative were correct, we would expect there to be some tangible evidence of an older metamorphic or igneous event in these rocks. To our knowledge, no intrusives older than Cambrian in age have been recognized in these specific rocks, nor is there petrographical evidence for an earlier highgrade metamorphic episode. The Olta metamorphic complex, in the southern Sierras of La Rioja province, is somewhat problematical in that no metamorphic ages were obtained and the maximum detrital age spans the Neoproterozoic-Cambrian boundary (Fig. 6b). Dahlquist & Baldo (1996) have suggested that anatexis in this complex was related to high heat-flow associated with granitoid emplacement (Chepes igneous complex of Pieters et al. 1997) during the Ordovician. However, both Dahlquist & Baldo (1996), and Pieters et al. (1997) recognized an earlier, low-grade, deformational episode in these metasedimentary rocks. Pankhurst et al. (this volume) suggest that this earlier episode occurred during the midCambrian. If this is so, the Olta metamorphic complex may be correlated with the Pichanas metamorphic complex and the Quilpo formation in the Sierras de Cbrdoba. Furthermore, an inferred late Neoproterozoic, maximum depositional age for these metasedimentary rocks suggests that they may be correlated with the late Neoproterozoic to early Cambrian, Puncoviscana Formation of Je~ek et al. (1985) and Willner & Miller (1986).

P a m p e a n orogeny

The young metamorphic rims on the detrital zircons in samples PL063 and PL147, as well as monazite growth in sample PL063 indicate that the basement rocks of the southern Sierras Pampeanas underwent high-grade metamorphism and deformation in the interval c. 540 Ma to c. 520Ma (Table 1). This time interval may

276

J. P. SIMS E T A L .

Fig. 11. Palaeozoic orogenic events in the southern Sierras Pampeanas. Metamorphic and intrusive events are shown on the right-hand side of the 'orogenic boxes', whereas tectonic events are shown on the left. See text for further discussion. be correlated with the Pampean orogeny of Acefiolaza & Miller (1982). High-grade metamorphism during the Pampean orogeny was of sufficient grade within the northern Sierras de C6rdoba to generate local partial melts (Baldo & Casquet 1996). These melts coalesced to form a number of highly

peraluminous granite bodies (Rapela et al. 1995; Rapela & Pankhurst 1996; Lyons et al. 1997). One of these bodies (El Pil6n Granite) has produced a Rb-Sr age of 523 + 4 M a (Rapela et al. this volume), and U - P b zircon ages of c. 527Ma (Lyons et al. 1997) and 523 + 2 M a (Rapela et al. this volume).

SIERRAS PAMPEANAS: PALAEOZOIC EVENTS One of the late calc-alkaline, I-type granites associated with the Pampean orogeny, which are abundant in the Sierra Norte de C6rdoba (Fig. 1), has yielded a magmatic age of 515 • 4 Ma (U-Pb, unpublished data). This represents a minimum age for the Pampean event. The significance of this age, which is younger than the collisional and metamorphic ages (see also Rapela et al. this volume), may be that the calcalkaline magmatism was not subduction-related as has been suggested by Ramos (1988), but rather related to post-collisional extension and melting of older I-type source rocks within the crust (e.g., Pitcher 1982; Chappell & Stephens 1988). Such an interpretation is consistent with the markedly felsic nature of the calc-alkaline granitoids in the Sierra Norte de Cdrdoba (e.g., Lira et al. 1997) compared with more typical, arc-related, I-type magmatism (e.g., Hildreth & Moorbath 1988; Silver & Chappell 1988). Furthermore, it would imply that the (Pampean) metamorphic basement, which predominantly consists of metasedimentary rocks, must be underlain by older crust (containing suitable I-type source rocks). This older crust is most likely to be the Rio de La Plata craton. Hence, it may be inferred that the Pampean orogeny involved overthrusting of a passive marginal basin eastwards over the continental crust. There is no compelling evidence at this stage, however, for subduction on the eastern margin of the Sierras Pampeanas as has been suggested by Ramos (1988).

Famatinian orogeny

Collision of Laurentia and Gondwana is generally considered to be responsible for the Ordovician Famatinian orogeny (e.g., Dalla Salda et al. 1992; Dalziel et al. 1994). However, MacNiocaill et al. (1997) argue from palaeomagnetic and faunal evidence that the two continents may never have collided; rather, that the Famatinian orogeny involved a complicated history involving the accretion of unique terranes, such as the Precordillera terrane, to the SW-Gondwana margin. If so, extensive development of a calc-alkaline, magmatic arc suggests that accretion of the Precordillera terrane was preceded by a period of eastward-dipping subduction on the Gondwana margin (e.g., Pankhurst et al. this volume). The detrital zircon data from JS129c also suggest that formation of this magmatic arc was preceded by or concurrent with a period of sedimentation that was predominantly sourced from the Pampean orogen. This sedimentary basin, which may have been

277

a back-arc basin, must have been incorporated into the collisional arc by the earliest Ordovician. The emplacement of mafic and ultramafic rocks of the Las Aguilas group at 478 + 6 M a (JSO8Of) and spatially associated orthogneiss at 484+ 7 Ma (JSO79e) constrains the age of the metamorphic peak in these Pampean-derived metasedimentary rocks to c. 480 Ma. This age is consistent with ages for the magmatic arc of between c. 495 Ma and c. 470 Ma (Pieters et al. 1997; Pankhurst et al. this volume). The time gap of c. 20 Ma between the latest Pampean intrusives (at c. 515 Ma) and the earliest Famatinian intrusives (at c. 495Ma), suggests that the two events may be linked. It could be inferred from the data and the geological relationships that the Pampean orogeny resulted from the initiation of compressional tectonics on the continental margin, whereas the Famatinian orogeny was due to the initiation of subduction and formation of a magmatic arc displaced to the west of the proto-continental margin. However, subduction was apparently shortlived and had probably ceased with the docking of Gondwana and the Precordillera terrane by c. 470Ma. The highest-grade metamorphic assemblages within the Pringles metamorphic complex, which developed at c. 480 Ma, are associated with westward-directed thrust fabrics (Sims et al. 1997). These compressional fabrics are overprinted by eastward-directed extensional fabrics containing distinctive tourmaline-apatite-bearing pegmatites which have produced c. 460U-Pb uraninite ages (Linares 1959). Furthermore, the Pringles metamorphic complex is structurally overlain by Ordovician sedimentary rocks of the San Luis Formation (Sims et al. 1997). These latter rocks were intruded by tonalite and granodiorite at c. 470 Ma (Sims et al. 1997; Fig. 4) yet neither contain the distinctive pegmatites nor show evidence of Famatinian-age structural fabrics. The monazite and high Th/U zircons from the Pringles metamorphic complex (JS129c), however, have given younger ages of c. 450 Ma. This disparity in ages suggest a more complex basement-cover relationship than is immediately apparent. It suggests that the basement (the Pringles metamorphic complex) and cover (the San Luis Formation), were not contiguous during the Famatinian orogeny, but must have been juxtaposed at a later stage. Evidence for cooling below c. 350~ (the closure temperature for the Ar-Ar isotopic system in muscovite; McDougall & Harrison 1988) and waning of the Famatinian orogeny in the basement terrane at c.450Ma is provided by Ar-Ar data from muscovite in a shear zone in

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the Sierra de Chepes (PP192b; Table 2). This age also corresponds well with Rb-Sr age data of Pankhurst et al. (this volume), from various intrusive phases of the Chepes igneous complex in the southern sierras of La Rioja Province, which suggest that the collision-related granitoid magmatism had ceased by the mid-Caradoc (c. 450 Ma).

a discrete Devonian event within the southern Sierras Pampeanas. This differs significantly from previous interpretations (e.g., yon Gosen & Prozzi 1996) that incorporated this deformation episode within the Famatinian orogeny, related to collision of the Precordilleran terrane.

Conclusions Achalian orogeny The Ar-Ar data for the low-grade shear-zones in the southern Sierra Pampeanas have produced dominantly Devonian ages (Table 2). In the sample from the Guamanes Shear Zone in the Sierras de C6rdoba (PL230b), biotite defining the shear fabric produced a well-defined age plateau of 358 4-2 Ma (Fig. 10f), whereas muscovite from an earlier assemblage in the same rock was only partially reset by the shearing event and produced a continuously rising age spectra (Fig. 10e). The age for the Guamanes Shear Zone is consistent with the Ar-Ar age from the Rio Guzman Shear Zone in the Sierras de San Luis (JS13Ob) that produced a rising age spectra in the range of 351-362 Ma. In contrast, the shear-zone sample containing the kyanitebearing assemblage in the Sierras de San Luis (ACOllb) produced a slightly older age of 366 42Ma, and the strike-slip, shear-zone sample (JSO86f) produced a plateau age of 375 + 1 Ma. All these shear fabrics overprint granites with early Devonian U-Pb zircon ages. For example, the Escalerilla Granite in the Sierras de San Luis has produced a U-Pb zircon age of 403 + 6 Ma (Sims et al. 1997; Fig. 4). This constrains the maximum age of shearing to the early Devonian. Astini (1996) suggests that within the Precordillera, a significant compressional event occurred due to collision of the Chilenia terrane at the Siluro-Devonian boundary and termed the event the 'Fase Precordiller~inica'. Our data suggest that the effects of this collisional event extended beyond the western Sierras Pampeanas as suggested by Astini (1996) and is manifest in the form of transpressional shear-zones in the southern Sierras Pampeanas. These shear-zones record components of both thrust and sinistral displacement that continued to the beginning of the Carboniferous. This contrasts with the 'Fase Precordillerfinica' of Astini (1996) in which the Precordilleran terrane was effectively undergoing post-collisional extension for most of the Devonian. Thus the Achalian orogeny may be correlated with the compressional phase of the 'Fase Precordillerfinica' of Astini (1996) and defined as

New geochronological data obtained from the southern Sierras Pampeanas place important constraints on the Palaeozoic tectonic development of the western-Gondwana margin. The following conclusions may be drawn from the data. (1) The oldest metasedimentary basement rocks in the southern Sierras Pampeanas have late Neoproterozoic maximum detrital ages and may be as young as early Cambrian in age. (2) Initial deformation and high-grade metamorphism of the metasedimentary rocks occurred during the Pampean orogeny in the interval c. 540 to c. 520 Ma. (3) A period of sedimentation occurred on the Gondwana margin following or closely associated with the Pampean orogeny that was largely sourced from Pampean-age rocks. (4) Sedimentation had ceased by c. 480Ma when mafic and ultramafic rocks, and felsic orthogneiss were emplaced in a collisional setting associated with the Famatinian orogeny. (5) Extensional collapse of the Famatinian orogenic belt occurred during the period c. 470 Ma to c. 450 Ma and was followed by a period of relative quiescence in the southern Sierras Pampeanas terrane for approximately 50 Ma. (6) Re-initiation of convergence on the Gondwana margin resulted in dispersed lowgrade shearing associated with the Achalian orogeny that extended for the duration of the Devonian. This study was conducted under the auspices of the Geoscientific Mapping of the Sierras Pampeanas Cooperative Project between the Australian Geological Survey Organisation (AGSO) and the Servicio Geol6gico Minero Argentino (SEGEMAR). We would like to thank all those who contributed to the running of the project or participated in the field programmes. In particular, Roberto Page and Lynton Jaques are acknowledged for their continued support of the project. Graeme Rogers and an anonymous referee are thanked for their constructive reviews. George Gibson, Richard Blewett and Lynton Jaques are also thanked for their comments on an earlier version of the manuscript. This paper is published with the permission of the directors of AGSO and SEGEMAR.

SIERRAS P A M P E A N A S : P A L A E O Z O I C EVENTS

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MARTINO, R. D. 1993. La faja de deformaci6n Guamanes: petrografia, estructura interna y significado tect6nico, Sierra Grande de C6rdoba, Argentina. Revista de la AsociaciOn Geol6gica, Argentina, 48, 21-32. McDOUGALL, I. & HARRISON, T. M, 1988. Geochronology and Thermochronology by the 4~ Method. Oxford University Press, New York. -& ROKSANDIC, Z. 1974. Total 4~ fusion ages using HIFAR reactor. Journal of the Geological Society of Australia, 21, 81-89. MUIR, R. J., IRELAND, T. R., WEAVER, S, D. & BRADSHAW, J. D. 1996. Ion microprobe dating of Paleozoic granitoids: Devonian magmatism in New Zealand and correlations with Australia and Antarctica. Chemical Geology (Isotope Geoscience), 127 191-210. ORTIZ SUAREZ, A., PROZZI, C. & LLAMBIAS,E. J. 1992. Geologla de la parte sur de la Sierra de San Luis y granitoides asociados, Argentina. Estudios Geologicos, 48, 269-277. PA~ES, J. B. & MILLER, J. D. 1989. Precise U-Pb ages of Duluth Complex and related mafic intrusions, Northeastern Minnesota: Geochronological insights to physical, petrogenetic, paleomagnetic, and tectonomagmatic processes associated with the 1.l Ga midcontinent rift system. Journal of Geophysical Research, 98, 13 997--14 013. PANKHURST, R. J., RAPELA, C. W., SAAVEDRA, J., BALDO, E., DAHLQUIST, J. & PASCUA, i. 1996.

Sierras de Los Llanos, Malanzfin and Chepes: Ordovician I and S-Type granitic magmatism in the Famatinian Orogen. XIII Congreso Geol6gico At2gentino y III Congreso de Exploracidn de Hidrocarburos, Buenos Aires, Actas, V, 415-416. & FANNING, C. M. i 998. ~I'he Famatinian magmatic arc in the central Sierras Pampeanas. This volume. PIETERS, P., SKIRROW,R. G. & LYONS,P. 1997. Informe geol@ico y metalogdnico de las Sierras de Las Minas, Chepes y Los Llanos (provincia de La Rioja), 1: 250,000. Anales XXVI, Instituto de Geologla y Recursos Minerales, SEGEMAR, Buenos Aires. PITCHER, W. S. 1982. Granite Type and Tectonic Environment. In: Hs~?, K. J. (ed.) Mountain Building Processes. Academic Press, London, 000-000. RAMOS, V. A. 1988. Late Proterozoic - early Paleozoic of South A m e r i c a - a collisional history. Episodes, 11, 168 174. RAPELA, C. W. & PANKHURST, R. J. 1996. The Cambrian plutonism of the Sierras de Cdrdoba: pre-Famatinian subduction? and crustal melting. ,VIII Congreso Geol6gico Argentino y Ill Congreso de Exploraci6n de Hidrocarburo, Buenos Aires, Actas, V, 491. , BALDO, E. & SAAVEDRA,J. 1995. Cordierites in S-Type granites: restites following low pressure, high degree partial melting of metapelites. The Origin of Granites and Related Rocks. US Geological Survey Circular, 1129, 120 121. , CASQUET, C., BALDO, E., SAAVEDRA, J., GALINDO, C. & FANNING, C. M. 1998. The Pampean Orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de C6rdoba. This volume. , TOSELLI,A., HEAMAN, L. & SAAVEDRA,J. 1990. Granite plutonism of the Sierras Pampeanas; An inner cordilleran Paleozoic arc in the southern Andes. In: KAY, S. M. & RAPELA, C. W. (eds) Plutonism from Antarctica to Alaska. Geological Society of America Special Papers, 241, 77-90. SAMBRIDGE, M. S. & COMPSTON, W. 1994. Mixture modelling of multi-component data sets with application to ion-probe zircon ages. Earth and Planetao' Science Letters, 128, 373-390. SILVER, L. T. ~; CHAPPELL,B. W. 1988. The Peninsular Ranges Batholith: an insight into the evolution of the Cordilleran batholiths of southwestern North America. Transaction of the Royal Society of Edinburgh." Earth Sciences, 79, 105-121. SIMS, J. P., SKIRROW, R. G., STUART-SMITH, P. G. & LYONS, P. 1997. Informe geol6gico y metalog~nico de las Sierras de San Luis y Comechingones (provincias de San Luis y C6rdoba), 1:250,000. Anales XXVIII, Instituto de Geologla y Recursos Minerales, SEGEMAR, Buenos Aires. STEIGER, R. & J~,GER, E. 1977. Subcommission on geochronology: Convention on the use of decay constants in geo- and cosmochronology. Earth and Planetary Science Letters, 36, 359-362. TETLEY, N. I., McDOUGALL, I. & HEYDEGGER, H. R. 1980. Thermal neutron interferences in the 4~ dating technique. Journal of Geophysical Research, 85, 7201 7205.

SIERRAS P A M P E A N A S : P A L A E O Z O I C EVENTS UNRUG, R. 1992. The supercontinent cycle and Gondwanaland assembly: component cratons and the tinting of suturing events. Journal of Geodynamics, 16, 215-240. YON GOSEN, W. & PROZZI, C. 1996. Geology, structure, and metamorphism in the area south of La Carolina (Sierra de San Luis, Argentina).

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granitic rocks. Transactions of the Royal Society of Edinburgh: Earth Sciences, 83, 447-458. ---, CHAPPELL, B. W., MCCULLOCH, M. T. & CROOK, K. A. W. 1991. Inherited and detrital zircon - Clues to the early growth of crust in the Lachlan Fold Belt. Geological Society of Australia Abstracts, 29, p. 58 WILLYER, A. P. & M1LLER, H. 1986. Structural division and evolution of the lower Paleozoic basement in the NW Argentine Andes. Zen-

tralblatt fur Geologie und Paldontologie, I, 1245 1255.

The Early Palaeozoic magmatic record of the Famatina System: a review J. S A A V E D R A 1, A. T O S E L L I 2'3, J. R O S S I 2, E. P E L L I T E R O 4 & F. D U R A N D

2,3

1Instituto de Recursos Naturales y Agrobiologia, CSIC, Apartado 257, 37080 Salamanca, Spain 2 Facultad de Ciencias Naturales e Instituto Miguel Lillo, UNT, Miguel Lillo, 205, 4000 Tucum(m, Argentina 3 Instituto Superior de Correlaci6n Geol6gica, CONICET, Miguel Lillo, 205, 4000 Tucum6n, Argentina 4 Departamento de Geologia, Facultad de Ciencias, Universidad de Salamanca, 37071 Salamanca, Spain Abstract: The Famatina System is an elongated geomorphological zone between the Eastern and Western Sierras Pampeanas, characterized by important Ordovician volcanism. The basement consists mainly of a sedimentary protolith, metamorphosed to low grade during the Pampean orogeny (Early-Mid-Cambrian), together with sparse outcrops of highgrade metamorphic rocks. Overlying this basement, the Volcancito Formation consists of thin sediments of uppermost Cambrian and Tremadocian age, deposited in restricted basins. A much more voluminous series of volcanic and sedimentary deposits, initiated in the Arenig and culminating in Early-Mid-Llanvirn times, indicates a significant increase in the rate of tectonic subsidence. The form of these deposits shows that they were laid down in an ensialic marginal basin that developed contemporaneously with active volcanic activity. The products of the latter were predominantly acidic, generated by crustal anatexis, with much rarer occurrence of basaltic and trachyandesitic compositions. The volcanic rocks are cut by granitoids of unknown absolute age, although the few radiometric ages obtained on them do not exceed 460 Ma. The volcano-sedimentary basin is interpreted as related to subduction, and the granitic and volcanic rocks together represent a continental magmatic arc developed regionally during the Early Ordovician.

The Famatina System (Sistema de Famatina) of NW Argentina lies between the Eastern and Western branches of the Sierras Pampeanas (Fig. 1). It exhibits intense magmatism developed within the Famatinian Cycle (see Pankhurst et al. this volume), which began in the late MidCambrian (Acefiolaza & Toselli 1976). Its main characteristic is the marked presence of Ordovician volcanism. Another important volcanic event occurred in the Tertiary, as well as further less important events of doubtful age. Recently published syntheses deal with the geology in general (Acefiolaza et al. 1996a), or certain specific petrological aspects of the Lower Palaeozoic rocks (Rapela et al. i992; Toselli et al. 1996a). The present work addresses the relationships between the Early Palaeozoic magmatism and the regional metamorphic rocks, the characterization of the volcanic rocks of this age, and the interpretation of the Ordovician sequences of the Famatina System, which are the host rocks of many

granitoids. A geodynamic model to account for the development of the Famatina System is proposed. Relationships in the immediately surrounding geological settings are considered in other studies in this volume. We use the time scale proposed by Gradstein & Ogg (1996).

Metamorphic basement The basement is heterogeneous in both lithology and age. The outcrops with the highest metamorphic grade are also the oldest ones and are located in the western border of the Famatina System (Fig. 1). These rocks constitute the socalled Espinal Formation of the Western Sierras Pampeanas and are mainly composed of schists, gneisses, amphibolites and migmatites. Radiometric dates from a number of different sites, mostly to the south of the Famatina System itself, have shown that the high-grade metamorphic

SAAVEDRA,J., TOSELI, A., ROSSI,J., PELLITERO,E. & DURAND, F. 1998. The Early Palaeozoic magmatic record of the Famatina System: a review. In: PANKIJURST,R. J. & RAPEtA, C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 283-295.

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Fig. 1. Geological map of the Famatina system (modified from Acefiolaza et al. 1996a)

MAGMATIC RECORD OF THE FAMATINA SYSTEM event may be assigned to the Grenville Orogeny, with ages around 1000 Ma (Varela & Dalla Salda 1992; McDonough et al. 1993; Varela et al. 1996). Within the Famatina System, the basement includes migmatitic gneisses with mineral parageneses corresponding to aluminous rocks in the southern part of the Sierra de Paganzo (Saa11993) and amphibolites and granitic orthogneisses in the Sierra de Copacabana (L6pez et al. 1996). The high-grade rocks are cut by granites, and basic bodies are also known in these two instances. However, the most widespread development of the basement of the Famatina System consists of sedimentary rocks that were metamorphosed to low grade during the Neoproterozoic-Early Cambrian Pampean Cycle (approximately equivalent to the Brazilian and Pan-African cycles; Acefiolaza & Toselli 1976). Durand et al. (1992a) have divided this basement into the La Aguadita and Negro Peinado formations, and consider both to be representative of Pampean sedimentation west of the Rio de La Plata craton. In the western part of the Sierra de Paim/m (Fig. 1), the former is defined by the presence of Ptanolites sp. and Paliella sp., whereas the Negro Peinado Formation is unfossiliferous. Sedimentation in the corresponding basin (a passive margin marine sequence, with shelf carbonates and deep-water turbidites) probably occurred partly on oceanic crust (Rapela et al. 1992). Sedimentological studies of both formations (Clemens 1993; Rossi et al. 1996) indicate sequences with a predominance of fine-grained sediments, varying in thickness, and with features of low-density turbidite transport. The metagraywackes comprise quartz, feldspars and lithic materials with a chlorite and sericite matrix; zircon, anatase, tourmaline, apatite, ilmenite and pyrite are present as accessories. Rossi et al. (1996) envisaged provenance from granitic to quartz-monzonitic sources on a passive continental margin. Some volcanic or volcano-sedimentary intercalations occur; these are acidic, poorly developed and very localised. In terms of Pampean deformation, the metapelites display a subvertical S1 parallel to SO, and a cross-cutting $2 foliation.

Latest Cambrian to Early Ordovician sedimentation In the north of the Famatina System (at latitudes less than 29~ approximately that of Villa Uni6n) there are Lower Palaeozoic outcrops of varying dimensions, the smallest of which cannot be mapped at the scale of Fig. 1.

285

The basal unit is the Volcancito Formation, for which Toselli (1975) established a Lower Tremadoc age on the basis of a graptolite fauna (accompanied by trilobites and brachiopods). In this sequence, affected by contact metamorphism, volcanic rocks are either absent or of doubtful age. Previous studies (Acefiolaza et al. 1996a, b, and references therein) have shown that this formation mainly comprises black lutites, typical of anoxic environments, dark sandstones and subordinate limestones, with evidence of sedimentation below wave level on an external shelf. The proximal facies, which includes carbonate levels, contains trilobites and conodonts with ages ranging from Uppermost Cambrian to the basal Arenig (?), while the distal ones, with graptolites, are Tremadocian. The presence of lithoclasts of lutites and phyllites from the La Aguadita and Negro Peinado Formations indicates that these were the exposed in the source area. The same non-volcanic environment is inferred at a number of sites in this formation; overall, local variations can be seen in the palaeofauna, lithology, etc. In the southernmost outcrops there are fewer black lutites and few fossils (mainly trilobites) while in other zones graptolites predominate. In general, there are intercalations of millimetric laminae of black shale, typical of very low energy sedimentary environments, with massive beds of siltstones and mudstones indicative of turbidity currents, deposited in a restricted basin. The thickness of all these non-volcanic rocks is small, approximately 350 m.

Ordovician volcano-sedimentary sequences Overlying the sequences described above, there are elastic and volcaniclastic deposits (both locally fossil-bearing) as well as volcanic rocks of Arenig to Lower-Middle Llanvirnian age (Acefiolaza et al. 1996a, and references therein). Astini & Benedetto (1996) report that palaeocurrent directions indicate contribution of sediment to these deposits from both E and W. The lithological associations are very varied: conglomerates, sandstones, pelites, lavas, etc. Outcrop conditions do not permit clear observation of relations to the sedimentary series of the Volcancito Formation. These volcanic and continental terrigenous sediments are not observed to the south of 29~ Several different sedimentary environments, all related to active margin volcanism, have been identified in this sequence (Mfingano & Buatois

286

J. SAAVEDRA E T A L .

1996). The base is composed of fossiliferous mudstones, lava flows, volcanic breccias and conglomerates, indicating sedimentation in talus adjacent to a volcanic arc. Over these sediments occur finer-grained materials, sometimes fossiliferous: pelites, sandstones, siltstones and pyroclastic rocks, characteristic of shelf deposition close to volcanic centres. The top of the sequence consists of deltaic fan sandstones, pyroclastic rocks, breccias, volcaniclastic rocks and some lava flows suggesting a regime with weak volcanic activity. The thickness of this sequence exceeds 3500 m and, consequently, the rate of tectonic subsidence was much greater than for the Volcancito Formation. In the field, these rocks display significant metamorphism in the vicinity of later granites, in which they sometimes appear as xenoliths. In outcrop there is an alternation of porphyritic horizons with other of finer grain-size. Acidic compositions are predominant, especially dacites and rhyodacites; basaltic lavas are common but represent a relatively small volume. Plagioclase phenocrysts are completely altered, but retain evidence of strong zonation. Minerals such as biotite, pyroxene and amphibole are generally pseudomorphed by chlorite; the replacement of augite by amphibole and the development of secondary actinolite-tremolite is observable. In the most acid rocks (rhyolites) there are abundant graphic intergrowths of quartz and feldspar, with radial fibre or spherulitic textures. Quartz phenocrysts are notable either as idiomorphs or rounded and corroded by the melt, and needle-shaped zircon is a common accessory. Secondary minerals such as epidote and chlorite are of various types; some chloritization is contemporaneous with the epidote, but most is of later development. An important feature is the presence of piemontite, pumpellyite and prehnite. The ferromagnesian minerals are transformed to chlorite and opaque minerals more often than to calcite and titanite. The acidic rocks (rhyolites) are the least altered, but even in this group completely fresh samples are absent. In the basalts pyroxenes and amphiboles are common, but with an advanced degree of alteration. Augite has a typical hour-glass form but is largely transformed to amphibole and epidote. Serpentinized phenocrysts and zeolites are common. Plagioclase appears in three broadly defined grain-sizes: the phenocrysts are the most altered, those of intermediate size are the freshest, and the finest define miarolitic textures. Among the secondary minerals, chlorite (penninite) is much more abundant that titanite, epidote and calcite.

Other rocks show evidence of far more alteration. These are fine-grained and predominantly of microlithic texture, trachytic, with a groundmass of variably altered plagioclase, opaque minerals, calcite and chlorite. Sometimes pseudomorphs after olivine are seen, and at other times intergrowths of chlorite, actinolite and tremolite.

Nature of the volcanism As described above, the volcanic rocks have undergone intense alteration that has to a greater or lesser extent destroyed the magmatic minerals; significant primary remnants are exceptional. Compositions vary broadly from basaltic (subordinate) to rhyolitic; from moderately alkaline to subalkaline. ToseUi et al. (1990) demonstrated the bimodal nature of this volcanism and showed that major element compositions are not adequate to establish its classification. With the exception of a few rhyolites, all rocks have volatile contents >2%, so that it is necessary to revert to immobile element compositions. In Fig. 2a, a group of nine clearly basaltic members are distinguished from the more acid rocks; these coincide with those discriminated in the Ti-Zr plot (Fig. 2b). Other discrimination diagrams based on immobile element contents (Fig. 2c-f) show that they are mostly intraplate alkaline basalts, with a slight tendency to platemargin subalkaline basalt. Interestingly, no andesites s.s. appear, although slightly more alkaline equivalents (trachyandesites) are seen. Thi~blemont & T6gyey (1994) have proposed a diagram for the discrimination of volcanic and plutonic rocks with more than 55% SiO2 (Fig. 2g). This diagram is based on analyses of rocks classified according to well-known criteria and the fields defined include 90-95% of the samples in each case. For the Famatina basalts, almost all the data points fall in field B, that defined by rocks related to processes of crustal anatexis. The idea that they might belong to the continental tholeiite association can be ruled out since, according to these authors, partial fusion of average continental upper crust would lead to peraluminous magmas, while fusion of average greywacke would yield metaluminous and alkaline magmas. Mannheim & Miller (1996), using data from field, geochemical, mineralogical, petrological and isotopic data also showed the bimodal character if this volcanism and attributed the acid rocks to crustal anatexis at depths of 20-30 km. Overall it is suggested that the environment of volcanism was an ensialic basin that extended to

M A G M A T I C R E C O R D OF THE F A M A T I N A SYSTEM

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288

J. SAAVEDRA ET AL.

the N and S of the Famatina System (Bahlburg & Herv6 1997; Pankhurst et al. and Rapela et al. this volume, and references therein).

Plutonism The Lower Palaeozoic granitoids of the Famatina System have been divided according to their common characteristics into the following three groups: Cerro Toro, Cerro Blanco-San AgustinPaganzo, Narv~iez-lqufiorco-Safiogasta and, finally, Copacabana-Paim/m (see Figs 3 and 4) The Cerro Toro, Cerro Blanco, San Agustin and P a g a n z o plutons

These appear to the W and S of the Famatina System and may even be found in the eastern F o Copacabana na/~ A Paim~n

foothills of the Western Sierras Pampeanas (Toselli et al. 1988; Saavedra et al. 1992). The Espinal Formation, which is in contact with the Cerro Toro granite, displays a metamorphic paragenesis with garnet and kyanite, indicating pressures around 6kbar; this is similar to that given by hornblende geobarometry of the pluton (Rossi et al. 1991). There is clear and frequent evidence of assimilation of schists and gneisses, and ortho-amphibolite inclusions also occur. There is a marked foliation, without cataclasis, in these rocks. They range from granite to gabbros (Fig. 4), are coarse- to fine-grained equigranular, with biotite and hornblende, and exhibit evidence of magmatic interaction. The dominant evolutionary trend is calc-alkaline (Fig. 3), with aluminosity indices varying between 0.7 and 1.2. In the more basic members it is possible to observe tholeiitic features (Rapela et al. 1992). The Cerro

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MAGMATIC RECORD OF THE FAMATINA SYSTEM

289

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Fig. g. Geochemical patterns normalized to ocean ridge granites (ORG, Pearce et al. 1984) for representative granites from the Famatina System. Symbols as for Fig. 4.

Toro granite, which cuts the Las Planchadas Formation near Villa Castelli (Fig. 1), is a synto late-tectonic body; a Rb-Sr whole-rock isochron age of 4564-14Ma, with an initial 87Sr/86Sr ratio of 0.70967, has been reported by Saavedra et al. (1992).

In the Sierra de Paganzo, Saal (1993) recognised high temperature mylonite bands cutting the granites and passing transitionally into magmatic foliation, thus suggesting syn-kinematic intrusion in a ductile shear zone. To the south, a pressure of 3.0-3.5kbar (biotite-cordierite

290

J. SAAVEDRA E T AL.

geobarometer) was determined in the country rock gneisses, while the granite to the north and centre of the Sierra gave a pressure of 5.1 kbar (hornblende geobarometer). Rb-Sr ages of 450 + 7 Ma and 456 + 9 Ma were obtained, and 404 + 23 Ma in some minor granite bodies, with initial 8VSr/86Sr ratios between 0.706 and 0.709. The Cerro Blanco granites and their continuation (San Agustin), at the southwestern end of Famatina System (Fig. 1) crystallized at a pressure of 3.6kbar (Rossi et al. 1991).

The Narvdtez, N u ~ o r c o a n d Saffogasta plutons

Compared to the previous group, the chemistry of these rocks is clearly calc-alkaline, but much poorer in mafic members (Figs 3, 4). The Narvfiez granites and granodiorites (aluminosity index 0.8 to 1.7) have equigranular and porphyritic textures, with the development of granophyres and pegmatites. They contain accessory hornblende, epidote and biotite and their emplacement pressure was 2 kbar or less (Rossi et al. 1987). They contain few enclaves (of sedimentary and Ordovician volcanic rocks) and are covered by Carboniferous sedimentary rocks. The other plutons (aluminosity index 1.0-2.0) contain smaller tonalitic and syenogramtlc bodies as aplitic facies: they lack muscovite and contain allanite, epidote and titanite. There are numerous dykes of lamprophyre and granitic porphyry, as well as igneous enclaves, but pegmatites absent. The emplacement depth was between 2.0 and 3.5kbar (Toselli et al. 1996c) according to the paragenesis of the metamorphic aureole. Deformation has given rise to plastic flow, with rotation of feldspar. On the basis of a U-Pb zircon study, Loske & Miller (1996) give a tentative age of 459 Ma for these granites.

maximum age of 459Ma and initial 87Sr/86Sr ratios between 0.704 and 0.711. Deformation may have begun at very early stage, even before complete consolidation, in agreement with the dynamics of a broad regional extension regime (Durand & Ldpez 1996).

Petrogenetic and geochronological considerations Most of the granites are calc-alkaline, ranging from metaluminous to peraluminous (Rapela et al. 1992; Saavedra et al. 1992; Toselli et al. 1991, 1996a, b, and references therein). Their chemical variation is illustrated in Fig. 5, which shows patterns typical of volcanic arc or postcollisional granites. The total REE contents (58-197 ppm), the small or moderate Eu anomalies, and the La/Lu ratios (1.4-18.3) indicate early fractionation of plagioclase and, within the limitations expressed by Pearce (1996), active margin features. In the easternmost limit (Fig. 1), the Sierra de Fiambal also contains plutonic bodies older than those of the Famatina System (Grissom et aL this volume). Relationships with the country rocks always show that the granites are younger and clearly intrusive in the metamorphic or sedimentary envelope. Radiometric ages are lacking for most of the plutons; a few recent studies have given ages less than 460 Ma (see review of Acefiolaza et al. 1996a). Granites that form the southern prolongation of the Famatina System (the Los Llanos-Ulapes batholith) have similar petrological and geochemical features and Rb-Sr ages (Pankhurst et al. this volume), but the oldest granodiorite component is significantly older according to U-Pb zircon dating at around 490Ma (Pankhurst et al. this volume). Consequently it is not possible to discount the possibility that older granitoids might exist within the Famatina System.

The P a i m f n a n d C o p a c a b a n a p l u t o n s

Interpretation and discussion

This group of plutons is aligned NNW-SSE (Fig. 1). They are biotitic and two mica granites (aluminosity index 0.7-1.4) and are epizonal (pressures below 3 kbar, deduced from the mineralogy of the metamorphic aureole). In the Sierra de Paim~in they are intruded by gabbroid bodies, and bands of pegmatite alternate with the leucogranite (Durand et al. 1992b; Toselli et al. 1996d, and references therein). There are also phenomena of mingling with more basic rocks, some of which had shoshonitic affinities (Saavedra et al. 1996; P@rez 1991 and references therein). P6rez & Kawashita (1992) report a

Amalgamation of the various components of the basement of the Famatina System was largely accomplished during the Pampean orogeny, which affected a broad zone of the ProtoAndean margin from 17~ to the Sierras de C6rdoba at 33~ This was succeeded by an interval of some 30Ma of relative tranquillity before the start of Famatinian subduction (Rapela et al. and Pankhurst et al. this volume, and references therein). Following the Pampean orogeny in the Famatina System, which was manifested here by the low-grade metamorphism of the Negro Peinado and La Aguadita

MAGMATIC RECORD OF THE FAMATINA SYSTEM

291

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Fig. 6. Schematic cross-sections illustrating the proposed geodynamic model and the evolution of the Famatina System in the Lower Palaeozoic after the Pampean Orogeny (Middle Cambrian). (a) Volcancito Formation basin and equivalents, with proximal facies including thin layers of carbonates (Upper Cambrian- Lowermost Arenig) and distal facies (Tremadocian black lutites), forming a series a few hundred metres thick over the low-grade metamorphic rocks of the La Aguadita and Negro Peinado formations (Neoproterozoic-Lower Cambrian). (b) Genesis of the ensialic marginal basin of the Las Planchadas Formation and analogues rapid sedimentation of very thick (kilometre-range) volcanic and sedimentary materials (Upper Arenig-Llanvirnian), due to secondary extension caused by subduction. Thinning of the sialic crust, the subsequent ascent of mantle magma, and partial melting of the basement rocks (causing acid volcanism), result in the various volcanic products that characterise these overlapping processes. (e) Later emplacement of plutons, with clear features of crustal assimilation in favourable sites.

292

J. SAAVEDRA E T AL.

formations, this interval is represented in the Late Cambrian by the low tectonic subsidence rate of the Volcancito Formation (less than 400m of deposition from uppermost Cambrian to basal Arenig). Similarly, on a more regional scale, the Middle-Upper Cambrian sandstones of the Meson Group (NW Argentina-Bolivia) correspond to deposition on a slowly subsiding shelf at the border of an extensional basin, although is not clear whether this was a short-lived passive margin or a back-arc margin (Bahlburg & Herv~ 1997, and references therein). As in the Famatina, the subsidence rate increased markedly from Arenigian times onwards. In both cases, the thickness of sedimentary deposits during the Ordovician is some ten times that of a comparable time interval in the Cambrian. The Ordovician volcano-sedimentary sequences of the Famatina System, with their bimodal volcanism in ensialic marginal basins, show many similarities to deposits of this type from a wide range of geological ages (Tarney & Windley 1981), including Quaternary marginal basins of the Pacific region (Cole 1984). With respect to genetic classification, the fundamental points have been established in the syntheses of Tarney & Windley (1981) and Molnar & Atwater (1978); Uyeda and Kanamori (1979) classify subduction zones of 'Marianas' and 'Chilean' end-member types according to the ease of generation of back-arc basins (easy in the former and absent in the latter). Essentially, it is held that old ocean crust (>50 Ma), which is colder, denser and gravitationally more unstable than warm young crust, descends into the asthenosphere with greater ease and hence at a higher angle (60 ~ or more) in the 'Marianas' type; coupling between the upper and lower plate of the thin crust is weak and the creation of backarc basins is feasible. On continental margins these ensialic basins may range from purely sedimentary, to those showing the presence of basic volcanism, to those in which subcontinental crustal rupture occurred together with the appearance of oceanic crust. Ordovician subduction in the Famatina System is interpreted as an intermediate situation between the 'Marianas' and 'Chilean' endmembers of Uyeda & Kanamori (1979). The inferred crustal thickness, close to normal values of about 30 km, would account for the existence of a marginal basin; more marked thickening would have meant that the subduction-related extensional forces would have been insufficient to initiate back-arc opening. In such ensialic basins, it is possible to find evidence of more than one process. Subduction normally gives rise

to calc-alkalic magmatism, which can extend to more alkaline magmatism (as happens here) where extensional tectonics are accentuated by mantle upwelling. A similar case in Mexico has been reported by Nelson & Carmichael (1984) and features andesites, dacites and more acid members together with trachyandesites (benmoreites) and other alkaline rocks. The ascent of mantle in zones of crustal thinning may induce partial melting of the continental crust and generate acidic volcanic melts. This process is facilitated by the hydration that causes the descent of the oceanic slab, as has been reported for various areas (De Wit & Stern 1981; Cole 1984, and references therein). In the Famatina System these ensialic volcanic magmas were formed by anatexis of the metamorphic country rocks. Subsequent plutonic magmas were intruded into this series and part of the edge of the basin, even assimilating rocks from the western limit where significant amounts of basic bodies are found. At this time, in other parts of the Famatinian orogen, the crustal thickness or other factors affecting the genesis of these basins may not have allowed a similar development, the extensional effect being reduced to subsidence of sedimentary basins without volcanism. As has already been noted, the general features of the Famatinian plutonism (petrology, lithological associations, geochemistry, etc.) are continued in the Ordovician granitoid belt immediately to the south in the Los Llanos-Ulapes area, where they are emplaced into continental crust with Nd model ages of 1600 Ma (Pankhurst et al. this volume). Thus they are considered to be part of a single continental magmatic arc with subductionrelated plutonism and, in the Famatina System, related volcanism as well. Figure 6 shows a schematic view of the proposed evolutionary model. Sedimentation from the Upper Cambrian to the lower part of the Arenig is thin and differs from the later volcanicsedimentary sedimentation (Fig. 6a). Westward subduction (Fig. 6b) generated extension and rapid sedimentation of material derived from the exposed metamorphic basement, together with contemporaneous volcanic material, on the continental margin. This did not, however, lead to complete rupture with generation of newly formed ocean floor. At the same time, there was a mantle doming effect that led to melting of the sialic crust and generation of acid volcanism (Fig. 6b). Subsequent plutonism (Fig. 6c) was generally calc-alkalic and heterogeneous, with clear local evidence of assimilation of the crustal wall rocks, including those of the Western Sierras Pampeanas.

MAGMATIC RECORD OF THE FAMATINA SYSTEM

Conclusion The basement rocks with the greatest areal distribution in the Famatina System are those of the Negro Peinado and La Aguadita formations, which have a low-grade metamorphism developed during the Pampean Orogeny. Over these was laid a thin sedimentary series, the Volcancito Formation, from the Upper Cambrian to the Tremadoc-Arenig boundary, deposited in a restricted basin. The Famatinian orogeny was manifested by development, from the Lower Ordovician onwards, of a much deeper basin in which were deposited volcanic and sedimentary materials of diverse provenance during the Early and Mid-Ordovician. The majority of the volcanic rocks were acidic, some of moderate alkalinity and, in minor amount, basalts; the earliest magmas have features that indicate their origin by crustal anatexis. The characteristics are of a typical ensialic marginal basin, generated during subduction. There is very little geochronological data relating to the plutonic rocks, but those that exist suggest ages less than 460 Ma, younger than those recently reported by U - P b zircon dating of granitoids to the south of the zone. Nevertheless, similarities in the granitoids and associated rocks from both areas allow the identification of a single Famatinian continental arc developed from the beginning of the Ordovician. The authors gratefully acknowledge the comments of N. Petford and an anonymous reviewer, which contributed significantly to the improvement of the original manuscript.

References ACEIqOLAZA, F. G. & TOSELLI, A. 1976. Consideraciones estratigrfificas y tectdnicas sobre el Paleozoico Inferior del Noroeste argentino. H Congreso Latinoamericano de Geologia, Caracas, Venezuela, Actas, II, 755-763. - - - , MILLER, H. & TOSELLI, A. (eds) 1996a. Geologia del Sistema de Famatina. Mfinchner Geologische Hefte, 19A. Mtinchen, Germany. ACE)qOLAZA,G. F., TORTELLO,M. F. & ESTEBAN,S. 1996b. Oxygen deficient facies of Lower Paleozoic age in Northwestern Argentina. In: BALDIS,B. & ACENOLAZA,F. O. (eds), Early Paleozoic Evolution in N W Gondwana. Serie Correlaci6n GeoIdgica, Universidad Nacional de Tucum/m, Argentina, 12, 15-22. ASTINI, R. A. & BENEDETTO, J. L. 1996. Paleoenvironmental features and basin evolution of a complex volcanic arc region in the pre-Andean western Gondwana: the Famatina belt. III Inter-

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national Symposium on Andean Geodynamics, Saint Malo, France, R6sum6s ~tendus, 755-758 BAHLBURG, H. & HERVI~, F. 1997. Geodynamic evolution and tectonostratigraphic terranes of northwestern Argentina and northern Chile. Geological Society of America Bulletin, 109, 869-884. BOWDEN, P., BATCHELOR,R. A., CHAPPELL,B. W., DIDIER, J. & LAMEYKIE, J. 1984. Penological, geochemical and source criteria for the classification of granitic rocks: a discussion. Physics of the Earth and Planetary Interiors, 35, 1-1 I. CLEMENS, K. A. 1993. Sedimentologia, proveniencia y desarrollo geotect6nico del Sistema de Famatina en el Noroeste de Argentina. XII Congreso Geol6gico Argentino y II Congreso de Exploraci6n de Hidrocarburos, Mendoza, Actas, I, 310-320. COLE, J. W. 1984. Taupo-Rotorua Depression: an ensialic marginal basin of North Island, New Zealand. In: KOKELAR,B. P. & HOWELLS, M. F. (eds), Marginal Basin Geology. Volcanic and associated sedimentary and tectonic processes in modern and ancient marginal basins. Geological Society, London, Special Publications, 16, 109-120. DE WIT, M. J. & STERN,C. R. 1981. Variations in the degree of crustal extension during formation of a back-arc basin. Tectonophysics, 72, 229-260. DURAND, F. & LOPEZ, J. P. 1996. La deformaci6n dctil del flanco oriental del Sistema de Famatina. In: ACElqOLAZA, F. G., MILLER, H. & TOSELLI, A. (eds) q.v., 311-324. - - , ACElqOLAZA,F. G. & ESTEBAN, S. 1992a. El Proterozoico Superior y C~lnbrico Inferior metasedimentario de la regidn de Bordo Atravesado y extremo N de la Sierra de Paim/m, La Rioja, Argentina. In: ACEIqOLAZA,F. G. & ESTEBAN,S. (eds) El Paleozoico Inferior en Latinoam~rica y la G~nesis del Gondwana. Serie Correlaci6n Geoldgica, 9. Universidad Nacional de Tucum/m, Argentina, 31-40. - - , SAAVEDRA,J., PELLITERO,E., TOSELLI,A., RossI, J. & MEDINA, M. E. 1992b. The mafic-felsic complex of Anguinn (La Rioja, Argentina) and its signification into Andean low Paleozoic plutonism. Zentralblatt ffir Geologie und Palontologie, Stuttgart, Germany, Teil I, Heft 6, 1581-1595. GRADSTEIN, F. M. & OGG, J. G. 1996. A Phanerozoic Time Scale. Episodes, 19, 3-5. GRISSOM, G. C., DEBARI, S. M. & SNEE, L. W. 1988. Geology of the Sierra de Fiambal/t, northwest Argentina: implications for Early Palaeozoic Andean tectonics. This volume. IRVINE, T. N. & BARAGAR,W. R. A. 1971. A guide to the chemical classification of the common volcanic rocks. Canadian Journal of Earth Sciences, 8, 523-548. LOPEZ, J. P., IBAlqEZ PALACIOS, G. & TOSELLI, A. 1996. Geologia, petrografia y tect6nica de la Sierra de Copacabana, provincia de Catamarca, Argentina. XIII Congreso Geol6gico Argentino y III Congreso de Exploracion de Hidrocarburos, Buenos Aires, Actas, II, 171-177.

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LOSKE, W. & MILLER, H. 1996. Sistemtica U-Pb de circones del granito de lgqufiorco-Safiogasta. In: ACEIqOLAZA, F. G., MILLER, H. & TOSELLI, A. (eds) q.v., 221-228. MA,NGANO, M. G. & BUATOIS, L. A. 1996. Estratigrafia, sedimentologia y evoluci6n paleoambiental de la Formaci6n Suri en la subcuenca de Chaschuil, Ordovicico del Sistema de Famatina. In: ACElC,IOLAZA, F. G., MILLER, H. & TOSELLI, A. (eds) q.v., 51-75. MANNHEIM, R. & MILLER, H. 1996. Las rocas volcfinicas y subvolcfinicas eopaleozoicas del Sistema de Famatina. In: ACENOLAZA,F. G., MILLER, H. & TOSELLI, A. (eds), q.v., 159-186. MCDONOUGH, M. R., RAMOS, V. A., ISACHSEN,C. E., BOWRING, S. A. & VUJOVICH, G. I. 1993. Edades preliminares de circones del basamento de la Sierra de Pi6 de Palo, Sierras Pampeanas Occidentales de San Juan: sus implicancias para el supercontinente Proterozoico de Rodinia. XII Congreso Geol6gico Argentino y II Congreso de Exploraci6n de Hidrocarburos, Mendoza, Actas, III, 340-342. MESCHEDE, M. 1986. A method of discriminating between different types of mid-ocean ridge basalts and continental tholeiites with the Nb-Zr-Y diagram. Chemical Geology, 56, 207-218. MOLNAR, P. & ATWATER, T. 1978. lnterarc spreading and Cordilleran tectonics as alternates related to the age of subducted oceanic lithosphere. Earth and Planetary Science Letters, 41, 330-340. NELSON, S. A. & CARM1CHAEL,I. S. E. 1984. Pleistocene to recent alkalic volcanism in the region of Sanganguey volcano, Nayarit, Mexico. Contributions to Mineralogy and Petrology, 85, 321-335. PANKHURST, R. J., RAPELA, C. W., SAAVERDA, J., BALDO, E., DAHLQUIST,J., PASCUA,I. & FANNING, C. M. 1998. The Famatinian magmatic arc in the central Sierras Pampeanas. This volume. PEARCE, J. 1983. A "user's guide" to basalt discrimination diagrams. Unpublished Report, The Open University, Milton Keynes. - - 1 9 9 6 . Sources and settings of granitic rocks. Episodes, 19, 120-125. -& CANN, J. R. 1973. Tectonic setting of basic volcanic rocks determined using trace element analyses. Earth and Planetary Science Letters, 19, 290-300. & GALE, G. H. 1977 Identification of oredeposition environment from trace-element geochemistry of associated igneous host rocks. In: GASS, I. G. (ed.) Volcanic Processes in Ore Genesis. Geological Society, London, Special Publications, 7, 14-24. - - - , HARRIS, N. B. W. & TINDLE, A. G. 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of Petrology, 25, 956-983. PEREZ, A. 1991. Contribuci6n a la petrologia y geoqulmica de los granitoides y gabroides del N de la Sierra de PaimOn, Provincia de La Rioja. PhD thesis, Universidad Nacional de Tucumfin, Argentina.

& KAWASHITA, K. 1992. K/Ar and Rb/Sr geochronology of igneous rocks from the Sierra de Paim~in, northwestern Argentina. Journal of South American Earth Sciences, 5, 251-264. RAPELA, C. W., COIRA, B., TOSELLI, A. & SAAVEDRA, J. 1992. The Lower Paleozoic magmatism of Southwestern Gondwana and the evolution of the Famatinian orogene. International Geology Review, 34, 1081-1142. , PANKHURST, R. J., CASQUET, C., BALDO, E., SAAVERDA, J., GALINDO, C. & FANNING, C. M. 1998. The Pampean Orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de C6rdoba. This volume. RossI, J., DURAND, F. & TOSELLI, A. 1996. Aspectos estratigrfificos y geoquimicos del Basamento Metam6rfico del Sistema de Famatina, Argentina. XIII Congreso Geolbgico Argentino y III Congreso de Exploracion de Hidrocarburos, Buenos Aires, Actas, III, 679. - - - , TOSELLI,A., MEDINA,M. E. & SAAL,A. 1987. Los stocks granofiricos de Chaschuil, Sierra de Narv~iez, Catamarca, Argentina. X Congreso Geoldgico Argentino, Tucumfin, Actas, IV, 151-153. ~r WAGNER, S. 1991. Geobarometria de hornblendas en granitoides calcoalcalinos: Sistema de Famatina, Argentina. 7 ~ Congreso Geo16gico Chileno, Vifia del Mar, Actas, I, 244-247. SAAL,A. 1993. El Basamento Cristalino de la Sierra de Paganzo, La Rioja, Argentina. PhD thesis, Universidad Nacional de C6rdoba, Argentina. SAAVEDRA, J., PELLITERO, E., ROSSI, J. & TOSELLI, A. 1992. Magrnatic evolution of the Cerro Toro granite, a complex Ordovician pluton of Northwestern Argentina. Journal of South American Earth Sciences, 5, 21-32. & 1996. Magmatic affinities of iLowe'r Paleozoic) Granitoids of the Sierra de Paim/m, La Rioja, Argentina. XIII Congreso Geo16gico Argentino y Ill Congreso de Exploracion de Hidrocarburos, Buenos Aires, Actas, V, 467. TARNEY, J. & WINDLEY, B. F. 1981. Marginal basins through geological time. Philosophical Transactions of the Royal Society, London, A301, 217-232. THII~BLEMONT, D. & TI~GYEY, M. 1994. Geochemical discrimination of differentiated magmatic rocks attesting for the variable origin and tectonic setting of calc-alkaline magmas. Comptes Rendus de l'Academie des Sciences, Paris, 319 (s6rie I1), 87-94. TOSELLI, A., ROSSI, J., SAAVEDRA,J., PELLITERO,E. & MEDINA, M. E. 1988. Aspectos petrol6gicos y geoquimicos de los granitoides del entorno de Villa Castelli, Sierras Pampeanas Occidentales-Sistema de Famatina, Argentina. 5 ~ Congreso Geoldgico Chileno, Antofagasta, Actas, III, I17-28. - - - , DURAND, F., RossI, J. ~; SAAVEDRA,J. 1996a. Esquema de evoluci6n geotectfinica y magmtica Eopaleozoica del Sistema de Famatina y sectores de Sierras Pampeanas. XIII Congreso Geolbgico Argentino y III Congreso de Exploracion de Hidrocarburos, Buenos Aires, Actas, V, 443 -462.

M A G M A T I C R E C O R D OF T H E F A M A T I N A SYSTEM 8/; SIAL, A. N. 1996b. Granitos peraluminosos de la Zona Batolitica Central de Sierras Pampeanas (NW Argentino): Relaciones y significado geotectbnico. Memorias del XII Congreso Geol6gico de Bolivia, Tarija, II, 755-768. --, PELLITERO, E. SAAVEDRA, J., ROSSI, J. & MURATH, CH. 1996c. Granito Nufiorco-Safiogasta. In ACElqOLAZA, F. G., MILLER, H. & TOSELLI, A. (eds) Geologia del Sistema de Famatina. Mfinchner Geologische Hefte 19, Reihe A, Mtinchen, Germany, 211-219. - - , SAAVEDRA,J,. PELLITERO, E., ROSSI, J., ACEiqOLAZA, F. G. & MEDINA, M. E. 1990. Geoquimica y petrog+nesis del volcanismo Ordovicico de la Formacidn Las Planchadas, Sistema de Famatina. Revista de la Asociaci6n Geol6gica Argentina, 45, 313-322. , , MILLER, H. & ROSSI, J. 1996d. Interpretaci6n gen&ica de rocas plut6nicas de la Sierra de Paim~m, La Rioja, Argentina. Zentralblatt fr Geologic und Palontologie, Stuttgart, Germany, Teil I, Heft 7/8, 793-806. - - , WEBER-DIEFENBACH,K., ROSSI,J. & MILLER, H. 1991. Interpretacibn gen6tica de los granitoides del Paleozoico Inferior de Cerro Toro y lqufiorco, Sistema de Famatina, Argentina. 6 ~ Congreso Geol6gico Chileno, Vifia del Mar, Actas, I, 248252.

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Geology of the Sierra de Fiambalfi, northwestern Argentina: implications for Early Palaeozoic Andean tectonics GRADY

C. G R I S S O M t, S U S A N M. D E B A R I 2 & L A W R E N C E W. S N E E 3

1 Geology Deptartment, Stanford University, Stanford, California 94305, USA 2 Geology Department, Western Washington University, M S 9080, Bellingham, WA 98225, USA (e-mail: [email protected]) 3 US Geological Survey, Denver, Colorado, USA Abstract: Field mapping in conjunction with structural, metamorphic, and geochronological

data document the tectono-thermal history of exhumed deep crustal rocks in the Sierra de Fiambalfi, NW Argentina. The range consists of two structural blocks distinguished by different metasedimentary sequences and different grades of metamorphism. Orthogneiss and paragneiss in the northern structural block may have a Precambrian history. Greenschist- to amphibolite-facies metamorphism, intrusion, and injection migmatization affected all rocks at 540-550 Ma. A subsequent event in the Late Cambrian to Ordovician (e. 515 to 470 Ma) involved amphibolite- to granulite-faciesmetamorphism, mafic intrusion, and deformation, followed by cooling through mid-Palaeozoic time. The emplacement of Carboniferous (325-350 Ma) post-tectonic granites caused reheating and retrogression that was strongest toward the northeast part of the range. The Cambrian, Ordovician, and Carboniferous events in the Sierra de Fiambalfi were of regional extent as indicated by temporal correlations with events reported for other deep crustal rocks of the northern Sierras Pampeanas. Correlations between periods of intrusion and high-grade metamorphism in the northern Sierras Pampeanas and volcanic-sedimentary events in the adjacent supracrustal exposures confirm that rocks in the northern Sierras Pampeanas formed at deep (10-25 km) structural levels in the early Palaeozoic continental margin of Gondwana.

The foreland of the Andes in NW Argentina contains exposures of Palaeozoic rocks that formed and accumulated along the South American (Gondwana) continental margin. Late Precambrian to Palaeozoic low-grade metamorphic, sedimentary, and volcanic sequences crop out in the Puna Plateau, the foreland fold and thrust belt, and the Famatina System (Fig. 1). Farther south, Laramide-style reverse faults form ranges that expose medium- to high-grade metamorphic and intrusive rocks in a tectonic province called the Sierras Pampeanas (Fig. 1). The Palaeozoic sedimentary and volcanic sequences of the Puna Plateau, the foreland fold and thrust belt, and the Famatina System record Eocambrian passive margin sedimentation (Baldis & Omarini 1984; Je~ek & Miller 1987) followed by continental margin orogenesis during the early- to midPalaeozoic (Coira et al. 1982). The metamorphic and intrusive rocks of the Sierras Pampeanas have been interpreted as deep structural levels of the Palaeozoic continental margin sequence (Gonzfilez Bonorino 1950; Willner & Miller 1986; Willner et al. 1987). The medium- to highgrade metasedimentary rocks of the northern Sierras Pampeanas are inferred to be depositional equivalents of the Eocambrian passive margin sediments (Puncoviscana Formation) at

the base of the Palaeozoic sedimentary-volcanic sequences on the basis of lithological correlation (Acefiolaza & Toselli 1981), chemical correlation (Willner et al. 1985), temporal and palaeogeographical observations (Je~ek & Miller 1987; Je~ek et al. 1985), and structural mapping (Willner et al. 1987). If the metasedimentary rocks of the northern Sierras Pampeanas are indeed the depositional equivalents of the Puncoviscana Formation, their metamorphic-deformational record should reflect the same events recognized in those rocks and the overlying early Palaeozoic sedimentaryvolcanic section. However, a systematic comparison of the Late Precambrian to early Palaeozoic history of the sedimentary-volcanic rocks with that of the metamorphic and intrusive rocks in the northern Sierras Pampeanas has not been presented. Geochronological, structural, and metamorphic studies in the eastern portion of the northern Sierras Pampeanas (near Ancasti and Quilmes, Fig. 1) have delineated the timing and sequence of major metamorphic-deformational-intrusive events (Toselli & Rossi de Toselli 1973; Rapela 1976; Toselli et al. 1978; Miller & Willner 1981; Knfiver & Miller 1982; Rapela et al. 1982, 1990; Bachmann et al. 1986; Rossi de Toselli et al.

GRISSOM, G. C., DEBARI, S. M. & SNEE, L. W. 1998. Geology of the Sierra de Fiambalfi, northwestern Argentina: implications for Early Palaeozoic Andean tectonics. In: PANKHURST,R. J. & RAPELA,C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 297-323.

298

G. C. GRISSOM E T A L .

Fig. 1. Simplified geological map of NW Argentina. Neogene tectonic provinces, including the Puna Plateau, Fold and Thrust Belt, Famatina System, and northern Sierras Pampeanas, are bounded by heavy lines. The dotted-lines are international borders and the locations of the Sierra de Fiambal/t, the Sierra de Ancasti, and the Sierra de Quilmes are labelled. 1987; Willner et al. 1987). However, the timing and sequence of events farther west near the Sierra de Fiambalfi is not known. This paper presents field mapping, structural data, metam o r p h i c analysis, and geochronological data that d o c u m e n t the character and timing of multiple m e t a m o r p h i c - d e f o r m a t i o n a l - i n t r u s i v e events in the rocks of the Sierra de Fiamball_ located at the northwestern edge of the Sierras P a m p e a n a s (Fig. 1). The timing and character o f these events m a y be correlated with those of events d o c u m e n t e d near Quilmes and Ancasti (Fig. 1), allowing a systematic comparison with the early Palaeozoic sedimentary and volcanic events recognized in the c o n t e m p o r a neous supracrustal rocks of the P u n a Plateau, the foreland fold and thrust belt, and the F a m a t i n a System.

Analytical methods Mineral separates for U-Pb and 4~ dating were prepared by standard separation techniques and

hand-picked to greater than 99% purity. The details of analytical procedures and data collection for U-Pb ages are discussed in Wright & Fahan (1988). Radiogenic Pb was corrected for common Pb assuming 2~176 = 18.6 and 2~176 = 15.6. Sample dissolution and ion exchange chemistry were modified from Krogh (1973). All analysed zircon fractions were nonmagnetic at 1.7 A with a 20 ~ forward and 1~ side tilt on the Frantz magnetic separator. Uncertainties in the 2~176 and 2~176 measured ratios are on the order of 0.03-0.10%. Uncertainties in the 2~176 measured ratios range from about 0.3% to 13.0% depending on the ratio (all uncertainties are at two-sigma level). The uncertainties in the 2~176 ages were calculated on the combined uncertainties in mass spectrometry (principally the uncertainty in the 2~176 measured ratio) and an assumed uncertainty of • in the 2~176 ratio used for the common Pb correction. New data are presented in Table 1 and the new and unpublished results are summarized in Table 2. The gas extraction techniques and data collection for 4~ ages are discussed in detail in Snee et al. (1988). Plateau ages discussed in the text refer to the weighted average plateau criteria of Fleck et al. (1977), requiring that at least two adjacent heating steps,

GEOLOGY OF THE SIERRA DE FIAMBAL]t

299

o6 oo

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0 0 0

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0

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300

G. C. GRISSOM E T A L .

Table 2. U-Pb dating summal3, for Sierra de Fiambal6 Sample & mineral

Description

Location map

Interpreted age

Unnamed pluton Z13-zircon Unnamed pluton Z 11-zircon La Puntilla Zl-zircon Fiambalfi gabbro-norite Z 16 zircon Fiambal/t gabbro-norite Z16 monazite

Orthogneiss

Fig. 4a

Precambrian (crystallization)

Orthogneiss

Fig. 4a

540-550 Ma (crystallization)

Orthogneiss

Fig, 4b

c. 550 Ma (crystallization)

Tonalite in gabbro-norite

Fig. 9

Tonalite in gabbro-norite

Fig. 9

Migmatite leucosome Migmatite whole-rock

Fig. 9 Fig. 9

510-515 Ma (crystallization of Fiambalfi gabbro-norite) 505 Ma, 496 Ma, and 493 Ma cooling ages for different grain sizes 471 Ma cooling age 473 Ma cooling age

Granite

Fig. 13

>335 Ma (crystallization)

Granite

Fig. 13

c. 325 Ma (cooling)

Z15 monazite Ratones Z6 monazite Ratones Z4 zircon La Florida Z3 monazite

Ref.

References: 1, this study; 2, J. E. Wright unpublished data. comprising more than 50% of the total gas released, give the same age within uncertainty. Blocking temperatures of 325-270~ for muscovite (Cliff & Cohen 1980; Snee et al. 1988) and 300-260~ for biotite (Harrison et al. 1985; Snee 1982) are assumed in this study. Microcline spectra are interpreted assuming one or more blocking temperatures for different degassing domains between 150~ and 325~ (Harrison & McDougall 1981; Lovera et al. 1989). Additional analytical details and data are available as Supplementary Publication No. SUP 18119 (9 pp) from the British Library Document Supply Centre, Boston Spa, Wetherby, W Yorks, LS237BQ or the Geological Society Library. Electron microprobe analyses of minerals used for thermo-barometric calculations were performed using a JEOL 733 superprobe at Stanford University interfaced with a Kevex software system that uses the MAGIC 5 data reduction program. Accelerating voltages were 15kv, with a cup current of 15na. Counting times were 30 seconds and analytical errors are within • Standards included: quartz or albite for Si, kyanite, anorthite or spinel for A1, orthoclase for K, albite for Na, rutile for Ti, chromite for Cr, spessartine garnet for Mn, wollastonite for Ca, hematite for Fe, periclase for Mg, strontium fluoride for F, sodalite for C1, and zinc metal for Zn. The techniques are discussed in more detail in Grissom (1991, chapter 3).

Geological Outline of the Sierra de Fiambalfi The Sierra de Fiambahi consists of metamorphic and intrusive rocks thrust over Tertiary conglomerates and arkosic sandstones along a W-dipping reverse fault at the east side of the

range (Fig. 2). Within the range there are two major structural blocks referred to informally as the northern and southern blocks, separated from each other by the Ratones Granite and a N-S-trending reverse fault on the northwest side of the Ratones granite (Fig. 2). Metasedimentary units in the northern block consist of migmatitic paragneiss with interlayered quartzite and quartzo-feldspathic schist (Fig. 2). The paragneiss is interlayered with granitic orthogneiss on the east side of the northern block, whereas schist occurs to the west. Metasedimentary rocks in the southern block consist of Palaeozoic semipelitic schist, gneiss, and migmatite interlayered with calc-silicate schist and gneiss in a monoclinal sequence (Fig. 2). The data presented below document in chronological order the series of metamorphic, deformational, and intrusive events that are recorded by the rocks of the Sierra de Fiambalfi.

Precambrian An orthogneiss sample (Z13) from the northeast part of the Sierra de Fiambalfi yields a Precambrian intrusive age based on preliminary U - P b data on zircon (J. E. Wright pets. comm.). This may then imply a Precambrian metamorphic/ deformational history for the adjacent migmatitic country rocks, which display three generations of structures. From oldest to youngest the

GEOLOGY OF THE SIERRA DE FIAMBALA

301

to be Early Cambrian (see below). The foliation in the gneissic inclusions within leucosomes (Fig. 3) is defined by the alignment of biotite grains and ram-scale gneissic layering. The lenses of gneiss with a discordant foliation relative to the migmatitic layering may reflect the formation of Precambrian foliation and gneissic layering, prior to Cambrian injection migmatization.

Early Cambrian Northern block

Fig. 2. Simplified geological map of the Sierra de Fiambalfi. See location in Fig. 1. The locations of Figs 4 and 8 are shown. The northern and southern blocks refer to structural blocks described in the text.

Fig. 3. Field sketch of Precambrian paragneiss showing folded migmatitic layering with axial planar foliation (Sd). The sketch also illustrates gneissic inclusions in leucosome layers which display a discordant foliation and gneissic layering inferred to predate migmatization. The sketch was made approximately 1 km south of Location C in Fig. 4a. structures include a gneissic foliation, a migmatitic layering that overprints the gneissic foliation, and the dominant foliation (Sd) that overprints all of them (Fig. 3). The dominant foliation and folding event are Ordovician in age while the main leucocratic injections are inferred

Structural, metamorphic, and geochronological data related to Cambrian events in the northern block are summarized in Fig. 4a. The Cambrian event was characterized by injection migmatization during amphibolite-facies metamorphism in the southeast portion of the northern block (Fig. 4a) and the formation of a penetrative foliation during greenschist-facies metamorphism towards the northwest. The data will be discussed separately for the migmatitic rocks to the east and the schistose rocks to the west. The migmatitic layering in the eastern portion of the northern block is defined by leucosomes of quartz, oligoclase to andesine, and microcline, versus mesosomes (Johannes 1983) with quartz, biotite, andesine, ilmenite, and sometimes muscovite. The migmatitic layering is pervasive except within 1-500m wide quartzite layers which define sedimentary bedding. The migmatitic layering and the smaller quartzite layers were transposed parallel to the dominant foliation (Sd) by the youngest set of prominent folds (Fd, see description below). Sd cuts across migmatitic layering in the hinges of the youngest prominent folds (Fig. 3) indicating that the migmatitic layering is older than the dominant foliation. A single 5-15m wide layer of pelitic rocks observed within the quartzo-feldspathic migmatites at Location C (Fig. 4a) provides constraints on the grade of metamorphism during migmatization. The metamorphic assemblage (quartz, biotite, muscovite, plagioclase, garnet, tourmaline, ilmenite) and dominant foliation (Sd) in the pelitic rocks at Location C post-date migmatization. However, the pelitic rocks contain textural evidence for two generations of garnet growth. One textural variety of garnet contains abundant randomly orientated quartz inclusions and has embayed irregular margins; the other is inclusion-free with sub-idioblastic outlines (Fig. 5a). Both textural varieties of garnet form pressure shadows in the dominant foliation, indicating

302

G. C. GRISSOM E T AL.

Fig. 4. Summary of structural, metamorphic, and geochronological data related to Precambrian and Cambrian events in the northern block (a) and in the southern block (b). The map's location and legend are shown in Fig. 2. The grey area at the western edge of map (a) is part of the southern block. Diagonal lines in (a) denote the location of westward decrease in intensity of Early Cambrian migmatitic textures.

growth prior to, or synchronous with, the formation of the foliation. Electron microprobe analyses show that both varieties of garnet have the same rim composition but different core compositions (Fig. 5a), supporting the textural interpretation that the inclusion-rich garnets are older grains that re-equilibrated and/or were overgrown during a subsequent metamorphism. We suggest that the older inclusion-rich garnet is related to the migmatization event, although we cannot preclude the possibility that they are Precambrian in age. The absence of any high-temperature refractory phases such as sillimanite, spinel, or cordierite and the presence of muscovite in the pelitic rocks indicate that metamorphic temperatures did not exceed c. 700~ The age of migmatization is inferred from the intrusive relationships and crystallization age of granitoid intrusions in the northern block. Biotite+hornblende granodiorite and tonalite occur as tabular bodies 1-3 km in width, and as smaller lenses and pods meters to hundreds of metres across, within the migmatites. All bodies have a solid-state foliation parallel to Sd in the country rocks (see description below). Appendages emanate from the m-scale granitic lenses and layers to form the leucosomes in the country rock. The kilometre-scale intrusive layers have gradational contacts over 10-50 m. The contacts are characterized by increasing proportions of leucosome in the migmatized schist over a distance of 10-20 m, until wispy lenses of gneiss are enclosed in a matrix of homogeneous intrusive rock. The abundance of metasedimentary lenses then decreases over the next 20-30m into the intrusion, leaving massive foliated granodiorite or tonalite. The folding relationships (see below) and solid-state foliation require that the intrusions and injection migmatization predate the dominant foliation. The gradational contacts and appendages feeding into leucosomes indicate that the intrusions were synchronous with migmatization, and support the interpretation of Gonzlez Bonorino (1972) that the migmatites in the northern block were formed by lit-par-lit injection and are not the products of in-situ partial melting. A biotite orthogneiss (sample Z11; Fig 4a) yields a U-Pb zircon intrusion age of 540 to 550Ma intrusive age (J. E. Wright pers. comm.). Based on the gradational intrusive contacts with the migmatitic rocks, this is also interpreted as the age of injection migmatization and amphibolite-facies metamorphism in the eastern part of the northern block. Constraints on the depth of this Early Cambrian intrusion and migmatization can be derived

GEOLOGY OF THE SIERRA DE FIAMBAL/~

303

Fig. 5. Petrographical garnet textures and compositions in the northern block (a) and in the southern block (b). Early and late garnet refer to the texturally early and late varieties of garnet discussed in the text. Mineral textures were digitized from a projected image of the thin section. All thin sections were cut perpendicular to the dominant foliation. The trace of the dominant foliation is vertical. Full microprobe analyses are listed in the Appendix.

from mineral assemblages and the A1 content of amphiboles in the granitoid intrusions. The A1 content of igneous amphiboles has been calibrated as a geobarometer (Hammarstrom & Zen 1986; Hollister et al. 1987; Johnson & Rutherford 1989; Rutter et al. 1989). The igneous mineral assemblage in the intrusions includes quartz, plagioclase, microcline, biotite, titanite, ilmenite, -t-amphibole, +epidote, +apatite, and +zircon. Amphibole textures indicate early crystallization and in some cases surrounding clusters of biotite suggest resorption during late crystallization or subsequent metamorphism. The core and rim compositions of amphiboles in sample Z11 (Fig. 4a) are listed in Appendix A, and the pressures estimated for core and rim compositions are listed in Table 3. The foliated granitic plutons also contain subhedral to euhedral epidote which can be a liquidus phase in calcalkaline plutons at relatively high pressures of crystallization (Zen & Hammarstrom 1984). Epidote inclusions in biotite and hornblende

are consistent with the textural criteria of Zen & Hammarstrom (1984) for igneous epidote. Any interpretation of pressure determined from amphibole barometry is tentative, due to possible compositional changes in amphibole during metamorphism subsequent to crystallization. The homogeneous core and rim compositions of amphibole listed in the Appendix suggest

Table 3. Amphibole geobarometry of biotite orthogneiss Sample

Total AI*

P1

P2

P3

Z11 Rim Z11 Core

1.83 -t- 0.06 1.96 + 0.07

5.3 5.9

5.6 6.3

4.3 4.8

The sample location is shown in Fig. 4a. * Based on 23 oxygen formula, average for rim, with standard deviation. Calculated pressures (kbar): P1 (Hammarstrom & Zen 1986); P2 (Hollister et al. 1987); P3 (Johnson & Rutherford 1989).

G. C. GRISSOM ET AL.

304

(a)

- 10 c m

Mica-richbed Quartz-richbed Sd Bedding\ / r~.. x'~'~~x j Slin hinges

/ Bedding Sd Sample Location A, Fig. 4a (b)

- 2 cm S0 and S1

Crenulation lineation

M i c a - r i c h ~ bed Quartz-richbed

ce "",,S Oand S1

Sd/

Sample Location B, Fig. 4a (r

-1 m

Older Coarse-grained Folded Surfaceof /,'f ] ~ t Si)limaniteLineations Compositional~, / I Layering ( ~

Compositional\ ~\\~~ ' layenng , ~ \ ) Sd

Ld

Sample Location D, Fig. 4b Fig. 6. Block diagrams of textures and structures. (a) Fd fold hinge with axial planar Sd foliation defined by biotite and muscovite. Millimetre-scale sedimentary bedding defined by mica-rich and quartz-rich layers is folded, and remnant aligned grains of biotite and muscovite are related to S1. (b) Sd spaced cleavage that crenulates $1. Biotite, muscovite, and chlorite grow in the crenulation cleavage. The older penetrative foliation ($1) is defined by aligned biotite and muscovite. (e) Texturally early and late sillimanite. The cut face is perpendicular to Sd, folded compositional layering, and the texturally early sillimanite but is subparallel to the late lineated sillimanite.

that re-equilibration was either negligible or complete, since partial re-equilibration by continuous reactions would result in compositional zoning in the amphiboles. If the calculated pressures of 4-5 kbar reflect igneous conditions, they are consistent with field evidence indicating mesozonal to catazonal intrusion during amphibolite-facies metamorphism and injection migmatization. On the western side of the northern block, sedimentary bedding and two foliations are visible in the quartzo-feldspathic schist. Sedimentary layering is defined by millimetre- to centimetre-scale micaceous and quartz-rich layers and occasional 10-30m wide layers of pebble conglomerate and meta-sandstone. In some locations evidence of relict cross-bedding and graded bedding is visible in the pebble conglomerates and sandstones. The dominant foliation (Sd) visible in the migmatite can be traced continuously across the schist-migmatite contact (see below). Within the schist, Sd overprints an older foliation ($1). In the central part of Fig. 4a, near the contact with migmatite, $1 is only visible in the hinge region of folds, as remnant aligned grains of biotite and muscovite in mica-rich beds (Fig. 6a). Farther west the dominant foliation becomes weaker and $1 is visible in all beds as a penetrative foliation defined by the alignment of biotite, muscovite, and chlorite (Fig. 6b). S1 and sedimentary bedding are subparallel in the schist, approximately NNE-trending and steeply dipping (Figs 4a and 7a); dips are both to the east and west due to later folding. The metamorphic assemblage in the schist consists of quartz, biotite, plagioclase and muscovite in sandy layers and additional chlorite in mica-rich layers. Biotite, muscovite, and chlorite define $1 indicating greenschist-facies conditions during its formation. The age of Sj relative to the age of injection migmatization further east can be inferred from observations of the contact between them. Detailed mapping in two E-W traverses across the northern block shows that intensity of Early Cambrian injection migmatization dies out toward the northwest, in the strip denoted by the diagonal lines in Fig. 4a. Where observed, the change is transitional in nature. In the south, the gradation from migmatitic layering to a penetrative foliation (S0, both of which are overprinted by Sd, suggests that $1 in the schist formed during the Early Cambrian, synchronous with injection migmatization farther east. The correlation of $1 in the schist with migmatitic layering in the gneiss is important since it infers that the schist was not metamorphosed or deformed prior to the Early Cambrian.

GEOLOGY OF THE SIERRA DE FIAMBAL/i~

(a)

305

(b) Structural Domain: Northern Block N

/ /_

Structural Domain: Northern Block N

'.\ 9

\

9

o. I '~ 8.

\.

'

\ ~ Poles to S 0 9 Poles to S 1 9Fd fold hinges

". "~ ~ ,~7

, ', ~

o_ ,,;

%': {

(,,... + \'-,r

%.:..

")

i

f ~ z Crenulations on S1 9 re- Axis for S 1

Great Circle - Average Fd axial plane Dashed great circle - $1 girdle

* Poles to Sd Great Circle - Average Sd - Ld Stretching lineation

Fig. 7. Equal angle, lower hemisphere stereonets showing the orientations of structures in the northern block. So refers to sedimentary bedding visible in the schist on the western side of the northern block. $1 refers to the older penetrative foliation visible in the schist on the western side of the northern block. Sd, Ld, and Fd refer to the dominant foliation, lineation, and generation of folds visible throughout the northern block.

Southern block The data pertaining to Early Cambrian plutonism, metamorphism and deformation in the southern block is summarized in Fig. 4b. A crystallization age of c. 550 Ma may be estimated for the La Puntilla orthogneiss in the southwestern part of the southern block, based on the 2~176 ages of two discordant zircon fractions separated from sample Z 1 (data shown in Table 1). This is comparable with the age determined for the orthogneiss in the northern block, suggesting that the Early Cambrian history of the northern and southern blocks may have been similar. As in the northern block, younger metamorphism and deformation have strongly overprinted the Cambrian structures and mineral assemblages. The oldest visible deformational feature is compositional layering defined by varying mineral proportions in both calc-silicate rocks and sillimanite-bearing semipelitic rocks. This compositional layering is folded and transposed parallel to a younger, axial planar foliation that forms the dominant foliation in the southern block. The orientations of the compositional layering and sillimanite lineations measured on layered surfaces at fold hinges are shown in Fig. 4b. This coarse-grained, lineated sillimanite is only observed near the contact of the La Puntilla orthogneiss. It predates a finergrained sillimanite which grows in the dominant

foliation and occurs throughout the southern Sierra de Fiambalfi (Fig. 6c). In rare cases, relict biotite and/or muscovite are aligned parallel to the compositional layering. In addition to the relict sillimanite, there is also evidence for relict garnet in the southern block, with textures analogous to the older generation of garnet discussed for the northern block. The relict garnet occurs as large 1-2 mm grains with randomly orientated inclusions and irregular margins (Fig. 5b). A younger variety of garnet (< 1 mm) has sub-idioblastic outlines and no inclusions. As above, both textural varieties of garnet form pressure shadows for the dominant foliation. The relative ages of the two generations are indicated by overgrowths of inclusion-free garnet around inclusion-bearing cores. Both varieties have the same homogeneous compositions, with slight zoning < 10#m from grain margins (Fig. 5b). The homogeneous garnet compositions resulted from re-equilibration of all grains by volume diffusion at peak metamorphic conditions during the Late Cambrian to Ordovician (Grissom 1991, chapter 3). The randomly orientated inclusions of quartz, biotite, and plagioclase in the relict garnet may be indicative of the mineral assemblage with which the garnet originally equilibrated. The inclusion assemblage, in addition to the rare observation of relict muscovite aligned with the compositional layering, suggests that Cambrian metamorphism was of amphibolite-facies throughout

306

G. C. GR1SSOM ET AL.

the southern block. The spatial association of first generation sillimanite with the La Puntilla orthogneiss is interpreted as a local upgrade in Early Cambrian metamorphism, due to heat dissipated from the orthogneiss. Fig. 4b shows the sillimanite-in isograd, inferred to be of Early Cambrian age. The contact aureole interpretation is consistent with several observations. First, intrusion of the La Puntilla orthogneiss predates the dominant foliation, as does the first generation sillimanite in the adjacent country rocks. Second, the coarse-grained sillimanite occurs over a broad band (c. 2 km wide) adjacent to the orthogneiss, and there are no hornfelsic textures or other indications of a classic narrow contact aureole which pre-dated or post-dated the coarse sillimanite. Rather, the contact zone of the orthogneiss is characterized by an increase in the volume of leucocratic layers and a transition from schist to gneiss within 0.5kin of the contact. The occurrence of sillimanite in the thermal aureole of the orthogneiss provides P - T constraints for Early Cambrian metamorphism. The alumino-silicate phase relations of Holdaway (1971) indicate a minimum pressure of 2kbar and a maximum pressure of 8 kbar. Minimum temperatures near the La Puntilla orthogneiss are constrained by the alumino-silicate triple point to 500~ while maximum temperatures inferred from the presence of muscovite are approximately 700~ temperatures distant from the orthogneiss were probably lower. The ka Puntilla orthogneiss contains subhedral to euhedral epidote included in biotite and plagioclase. If any of the epidote is magmatic, its occurrence is consistent with the inferred mesozonal to catazonal emplacement level (Zen & Hammarstrom 1984). The deep contact aureole suggests that the growth of lineated sillimanite on the older compositional layering planes was synchronous with intrusion of the La Puntilla orthogneiss. Hence, the c. 550 Ma intrusive age of the orthogneiss constrains the age of this layering. The regional metamorphism and deformation which formed the compositional layering and older garnet throughout the southern block must therefore be Early Cambrian.

block are summarized in Fig. 8a. The dominant foliation (Sd) in the schist on the western side of the northern block is a sigmoidal crenulation cleavage that folds and crenulates the older S~. Biotite, muscovite, and sometimes chlorite grow in Sd crenulations (Fig. 6b). In this region, Sd is

Ordovician

Fig. 8. Summary of structural and metamorphic data related to the Late Cambrian to Ordovician event in the northern block (a) and the southern block (b). The map location and legend are shown in Fig. 2. Stretching lineation measurements marked with an S denote lineations defined by second generation sillimanite.

Northern block Structural data. The metamorphic and structural data related to Ordovician metamorphism, deformation, and intrusion in the northern

GEOLOGY OF THE SIERRA DE FIAMBAL,~ NNW-trending while sedimentary bedding and S~ are NNE-trending (Figs 7, 8). Farther east, but still within schist, bedding and S1 are transposed parallel to Sd. In this region Sd is a penetrative foliation defined by the alignment of biotite, muscovite, and rare chlorite (Fig. 6a). Sd is axial planar to tight, upright to moderately plunging folds (Fd) of sedimentary bedding and $1, with nearly N-S trending axial planes (Fig. 7a). Note that the hinges of metreto decametre-scale Fd folds are parallel to the crenulation lineation (microfold hinges) measured on $1 in the schists farther west (Fig. 7a). The dominant foliation can be traced continuously eastward into the migmatite on the eastern side of the northern Sierra de Fiambalfi (Fig. 8a). In the migmatite, Sd is defined by weak to strong alignment of biotite and sometimes flattened quartz grains. A weakly developed N-trending stretching lineation (Ld) is present on some Sd surfaces (Figs 7b, 8a). Ld is defined by the alignment of stretched quartz and biotite grains. Sd is axial planar to Fd folds and cuts across the folded migmatitic layering in Fd fold hinges (Fig. 3). Fd folds are tighter and more variable in orientation than in the schist farther west. Centimeter- and m-scale coaxial refolding of Sd and migmatitic layering by Fa+ folds is common and accounts for the east and west dips of Sd. Ordovician structures are also present in the older orthogneiss bodies of the northern block. The alignment of biotite and in some locations flattened quartz grains form a weak to moderate penetrative foliation (Sd). Sd in the intrusive rocks is parallel to So in the country rocks and to the intrusive contacts. Deformation post-dated crystallization, as evidenced by sub-grain formation in igneous quartz grains and recrystallization of igneous feldspar grains (1-2mm) at grain margins to grain sizes of 50-100#m. Pegmatitic lenses within the granites are commonly folded about axes parallel to Fo folds in the country rocks. The transition in S d from a spaced crenulation cleavage in the western schist to a penetrative foliation in the eastern migmatite reflects a west to east increase in Ordovician metamorphic grade. At the western margin of the northern block the metamorphic assemblage includes quartz, plagioclase, biotite, chlorite, and muscovite. Biotite, muscovite, and chlorite grow in the Sd spaced cleavage (Fig. 6b) indicating greenschist-facies conditions during deformation. Farther east the assemblage is the same except that chlorite is not present. Fig. 8a shows the distribution of samples which contain P - T conditions.

307

chlorite aligned with S d (labelled C) and samples without chlorite in Sd (labelled X). This distribution corresponds to the Ordovician chloriteout isograd. A narrow layer of semipelitic rocks at Location C (Fig. 8a) provides the best constraints on metamorphic grade in the migmatite on the east side of the northern block. The assemblage includes quartz, biotite, muscovite, oligoclase (An20-22), two generations of garnet, tourmaline, ilmenite, and retrograde chlorite. The Ordovician foliation (Sd) is defined by the alignment of biotite and muscovite. The younger, inclusionfree garnet forms clean contacts with biotite, muscovite, quartz, and oligoclase suggesting equilibrium between these phases. Garnet-biotite FeMg exchange thermometry and garnet, rutile, ilmenite, plagioclase, quartz (GRIPS) barometry (Bohlen & Liotta 1986) were performed on sample 655 from Location C (Fig. 8a). Experimental calibration of the endmember exchange reaction (Ferry & Spear 1978) was employed for garnet-biotite thermometry along with a garnet mixing model (Berman 1990) and a biotite mixing model (Indares & Martignole 1985). The Berman (1990) model was also employed for GRIPS barometry along with a plagioclase mixing model (Newton et al. 1980) and a mixing model for ilmenite (Stormer 1983); see Grissom (1991, Chapter 3) for additional details on the thermo-barometric formulations. The inclusion-free garnet (texturally late) from sample 655 shows slight core to rim zoning (Fig. 5a) which we interpret as growth zoning (Spear 1989). The garnet rim compositions are interpreted as reflecting peak metamorphic conditions during the Ordovician. Matrix biotite, ilmenite, and plagioclase grains yield relatively homogeneous compositions for different grains and within single grains. The results of garnet-biotite thermometry are shown in Table 4. Temperatures are listed for garnet rim compositions in conjunction with matrix biotite compositions. Note that the rims of both textural varieties of garnet give approximately the same temperatures, supporting the interpretation that first generation garnet crystals experienced new growth (or re-equilibration) at their margins during the growth of later garnet. These data suggest a peak metamorphic temperature of 620~ for Ordovician metamorphism at Location C (Fig. 8a). The results of GRIPS barometry are also shown in Table 4. Rutile was not observed in the mineral assemblage so all pressures calculated from the GRIPS barometer are maximum estimates. Note that the garnet rim compositions from both generations of garnet give

308

G. C. GRISSOM ET AL.

Table 4. Thermobarometric resultsfor sample 655* Textural variety of garnett

T(~ (Garnet-Biotite thermometry)

P (kbar) (GRIPS:~ barometry)

First generation core First generation rim Second generation core Second generation rim

-

< 5.4

614

<8.2

-

<7.0

623

< 8.12

* Sample 655 is from Location C, Fig. 8a. l-See text and Fig. 5 for a description of first and second generation garnet. SGRIPS, garnet, rutile, ilmenite, plagioclase, quartz barometer. All pressures are maximum estimates due to the absence of rutile in the barometric assemblage. See text for descriptions of the barometric and thermometric formulations employed.

the same pressures (c. 8 kbar), which is interpreted as the maximum pressure for Ordovician metamorphism. Timing of events. The maximum ages of Ordovician structures (S d and Fd) in the northern block are constrained by the intrusive ages of the orthogneiss bodies, which display welldeveloped Sd and metamorphic mineral growth. These give a maximum age of 540-550 Ma for deformation and metamorphism. The minimum age of Ordovician structures and metamorphism is constrained by 4~ step-heating ages on amphiboles separated from two samples in the northern block. The data are available as Supplementary Publication No. SUP 18119 (see p. 300) and the locations of samples 628 and 367 are shown in Fig. 9 along with the age spectra and isochron diagrams. Isochron ages and initial 4~ ratios were derived from the standard isochron plots of 4~ versus 39Ar/36Ar (Merrihue & Turner 1966). Error analysis is that of York (1966). Sample 628 is from a 50 m wide inclusion of mafic amphibolite within the paragneiss. The assemblage includes plagioclase, green amphibole, ilmenite and titanite. The alignment of amphibole grains and millimetre-scale layering of plagioclase and amphibole define a penetrative foliation parallel to Sd in the adjacent paragneiss. The step-heating spectrum shows unreasonably high ages for a small percentage of gas in the initial steps, and a relatively fiat spectrum between 390 and 410Ma for about

95% of the gas released. The linear array of heating steps 3 to 15 define an age of 395 + 4 Ma with minor excess 4~ resolved from the isochron analysis (4~ = 335 + 6). We interpret 3 9 5 + 4 M a to be the minimum age for effective closure of the amphibole to Ar diffusion. Metamorphic temperatures for sample 628 were between 450~ indicated by the amphibolite-facies assemblage, and 620~ estimated from geothermometry at Location C (Fig. 8a). Since effective closure temperatures for Ar diffusion in amphiboles are estimated to be 450-580~ (Harrison 1981), 395 Ma probably represents a cooling age and hence a minimum age for metamorphism. Sample 367 (Fig. 9) was collected within a post-deformational gabbroic intrusive body (northern gabbro) at the northeast corner of Fig. 8a. The igneous mineralogy of the gabbroic rocks consists of plagioclase, amphibole, clinopyroxene, ilmenite and biotite with no evidence of deformation or igneous foliation. The igneous assemblage is overprinted by static epidoteamphibolite-facies assemblages including hornblende after clinopyroxene, epidote after plagioclase, and titanite after ilmenite. Penck (1920) also interpreted the gabbro body as a postdeformational intrusion. This gabbro intrudes 540-550Ma orthogneiss layers, hence it has a maximum intrusive age of Early Cambrian. The step-heating spectrum for sample 367 shows initial high ages that decrease to a flat segment between 10 and 50% gas release (Fig. 9). From approximately 50 to 100% gas release (heating steps 4 to 16) the spectrum shows gradually increasing ages; steps 5 to 16 define a linear array on an isochron diagram with an apparent age of 406 + 3 Ma and slight excess 4~ (4~ = 302 + 26). The initial heating steps with high ages plot well above the linear array on the isochron diagram, suggesting the incorporation of excess 4~ Heating increments 7 and 8, which are labelled in Fig. 9, plot slightly below the best fit line on the isochron diagram. This may be attributable to excess argon; it cannot be attributed to high K impurities in amphiboles or degassing of two amphibole compositions with different K/Ca compositions because the ratio of K-derived 39Ar to Ca-derived 37Ar remains essentially constant through the last 95% of gas release. The step-heating spectrum is tentatively interpreted as excess argon (up to 50% gas release) overprinting a spectrum with decreasing ages attributable to argon loss (50 to 100% gas release). The isochron age of 4 0 6 + 3 M a is consistent with the 4~ plateau age for amphibole from sample 628.

GEOLOGY OF THE SIERRA DE FIAMBALA

309

Fig. 9. Sample locations and summary of the U-Pb and 4~ geochronological data related to Late Cambrian to Ordovician metamorphism. The black lines in the step-heating spectra indicate apparent age (left axis) and uncertainty. The dashed lines indicate 39Ar/37Arratios (right axis). The map is a simplified version of Fig. 2.

The effective closure of amphibole to Ar diffusion at > 4 0 0 M a in the northern block indicates that the intrusion of the northern gabbro, and hence the formation of the dominant foliation, must pre-date 400Ma. The 4~ amphibole ages, in conjunction with the crystallization age of the Cambrian pluton at Location ZI 1 (Fig. 4a), bracket the age of Fd

folds and the formation of Sd in the northern block between 400 and 550 Ma.

Southern block Ordovician metamorphism and deformation formed the prominent mineral assemblages and structures visible throughout the southern Sierra

310

G. C. GRISSOM E T AL.

(b) Structural Domain: Between the Fiambal~ (a) Structural Domain: Between the Fiambalfi gabbro-norite and Quebrada Cuestanzune (Fig. 8b) gabbro-norite and Quebrada Cuestanzune (Fig. 8b) N

.Fdhinges 9Poles to Sd \ Average Sd [] Ld stretching lineation o Long axis ofboudinaged mafic sills

/

+

9

~ 99 ~,~ 9

~..9 &

~rJ~ [] 7 /

(d) Structural Domain: Between the La Puntilla orthogneiss and Quebrada Cuestanzune (Fig. 6b)

hges 9 9

o,~

~ ~

"]

\ Average Sd X'~e~r " 9 Ld stretching lineation ~ _ (e)

\

\ \ Girdle to compositional layering

(C) Structural Domain: Between the La Puntilla orthogneiss and Quebrada Cuestanzune (Fig. 6b) N

9Poles to Sd

j ,/~-

[]PolestoFdaxialplanes ~ "o_* \ Avg Fd axial plane \o, ~ 9 Poles to compositional layering ~ 9 n - axis to compositional layering

(t)

Structural Domain: La Puntilla orthogneiss

9 Poles to Sd \AverageSd . ~9 Ld stretching lineation

/

/1/'/ J."

[] Poles to Fd axial planes \ Average Fd axial plane ~ ~%-9162 / o 9 ] 9 Poles tocompositional myenng ~ 9 / ~ / 9 n - axis to compositional layering * * 9 I~ 2~ 1/ 7 6 1 7 6 ,,o Early sillimanite lineation ~ / \ Girdle to compositional layering "--.-.___2.z/ Structural Domain: Fiambalfi gabbro-norite

N

"

N

t 9" 9Poles to Sd T~ 9 9 9 Fd hmges S 9 9 [] Poles to Fd axial planes ~O~' . "13 9 \ Average Fd axial plane o.S 9 o Poles to igneous layering (So) \ 9 % \ So girdle - - -~9 n - axis to So girdle

Fig. 10. Equal angle lower hemisphere stereonets showing structural data from the southern block. de FiambalL Figures 8b and 10 summarize the structural data and Fig. 11 summarizes the metamorphic data from Ordovician events in the southern block. Structural data. Sd is defined by the alignment of biotite, sillimanite, muscovite or hypersthene in semipelitic rocks, and biotite, amphibole, diopside, wollastonite, or titanite in calc-silicate rocks. It is concordant with the NNW-trending, steeply dipping monoclinal sequence of semipelitic and calc-silicate rocks. A NE-trending stretching lineation (Ld) defined by the alignment of biotite, sillimanite, sphene, amphibole, or stretched quartz grains is commonly visible on Sd surfaces, where it plunges nearly down-dip

(Figs 6c & 10). This is in contrast to the shallowplunging older lineation defined by coarse sillimanite and discussed earlier (Fig. 6c). Where Ld is defined by sillimanite, it is the texturallylater, finer-grained variety; these stretching lineations are marked with an S in Fig. 8b to illustrate that the later generation of sillimanite occurs throughout the southern block in contrast to the older sillimanite which only occurs near the La Puntilla orthogneiss (Fig. 4b). Sd is axial-planar to symmetric, isoclinal folds (Fd), which fold the Cambrian compositional layering described above. Fd folds occur on centimetre to decimetre scales, with gentle to steep plunges and symmetric geometry. Folding is confined within semipelitic and calc-silicate

GEOLOGY OF THE SIERRA DE FIAMBAL~k

Fig. 11. Summary of thermo-barometric and isograd data related to Late Cambrian to Ordovician metamorphism in the southern block. The geology corresponds to that in Fig. 2, but has been highly simplified for clarity.

layers and does not fold the NNW-trending monoclinal sequence of sediments (Fig. 2). Equal-angle stereo-net plots (Fig. 10a-d) show the relationships between Fd folds, folded compositional layering, and Sd for two structural domains in the southern block. In both structural domains, the r-axis defined by the girdled distribution of compositional layering coincides with the measured Fd hinges, and the average Sd is approximately parallel to the average Fct axial plane. Note that Fd axial planes and So are W-dipping in the domain adjacent to the Fiambal~i gabbro-norite and E-dipping in the domain adjacent to the La Puntilla orthogneiss. In both structural domains the orientation of Sd reflects the orientation of the nearest intrusive contact: the western contact of the Fiambalfi gabbronorite is vertical to steeply W-dipping, based on observations of vertical exposures as well as shallow (300m) drill holes (Villar & Segal 1988), whereas vertical exposures along the

311

eastern contact of the La Puntilla orthogneiss indicate that the contact is moderately to steeply E-dipping. Sd is concordant with both contacts and fans through vertical between the Fiambalfi gabbro-norite and the La Puntilla orthogneiss, suggesting that the feldspar- and pyroxenedominated intrusions acted as relatively rigid bodies in a more ductile quartz-dominated matrix of metasedimentary rocks. In the La Puntilla orthogneiss, the alignment of biotite, muscovite, and flattened quartz or feldspar grains define Sd (Fig. 8b). A steeply plunging stretching lineation (Ld) defined by stretched quartz and biotite grains is locally visible on Sd surfaces in the orthogneiss (Figs 8b, 10e). Pegmatitic dykes and pods within the orthogneiss are commonly folded with orientations coaxial to Fd folds in the metasedimentary rocks. Recrystallization of coarse (0.5-1.5 mm) igneous quartz and feldspar grains at grain margins to fine (50-100#m) sub-grains, and the growth of muscovite aligned with Sd after biotite, indicate that deformation and metamorphism of the orthogneiss post-dated crystallization. Rocks in the internal portions of the Fiambalfi gabbro-norite are commonly massive and unfoliated (DeBari 1994), but rocks at the margins of the gabbro-norite contain aligned amphibole, clinopyroxene, and orthopyroxene that define Sd. In the gabbro-norite, Sd is axial planar to folds of plagiogranite and sedimentary xenoliths. The orientation of folds in the gabbro-norite is coaxial with Fd folds in the country rocks (Fig. 10f). Igneous pyroxene grains (0.5-1.0 mm) with cumulate textures in the gabbro-norite are commonly recrystallized at grain margins to smaller (100-300 #m) aligned pyroxene or amphibole grains, indicating that deformation and metamorphism continued after crystallization. Discontinuous layers and lenses of mafic rock (1-500m across) in the metasedimentary rocks show similar mineral assemblages and structures those within the Fiambal~ gabbro-norite and are interpreted as small associated intrusions (DeBari 1994). Foliations defined by aligned pyroxene and amphibole are always concordant with the foliation in the adjacent country rocks. In some cases there is clear evidence for boudinage of a mafic layer to form lens-shaped inclusions. The long axes of the mafic boudins lie in the average Sd plane and are perpendicular to the La stretching lineations (Fig. 10a), suggesting that boudinage occurred during the formation of Sd. P-Tconditions. A detailed description of Ordovician metamorphism in the southern block is presented by Grissom (1991, chapter 3). The

312

G. C. GRISSOM ET AL.

results of isograd mapping and thermo-barometric calculations are summarized in Fig. 11. The muscovite-out isograd for semipelitic rocks, and the orthopyroxene-in isograd for mafic compositions, illustrate upgrading from amphibolite- to granulite-facies metamorphic conditions spatially associated with the Fiambal/t gabbro-norite. Additional isograds not shown in Fig. 11 include orthopyroxene-in for semipelitic rocks, sphene-out for mafic rocks, and amphibole-out, scapolite-in and wollastonite-in for calc-silicate rocks. A variety of geothermometers document a temperature gradient from approximately 700~ in rocks 2-3 kln from the Fiambalfi gabbro-norite to approximately 850~ at the gabbro-norite contact. These include garnet (rim)-biotite (matrix) temperatures of 720~ determined for sample 551, collected outside the Fiambalfi gabbro-norite aureole adjacent to the intrusive contact of the La Puntilla orthogneiss (Location F, Fig. 4b). The amphibolite-granulite transition is associated with the onset of muscovite and biotite dehydration melting, which formed migmatites in the vicinity of the Fiambalfi gabbro-norite. The pressure estimates from a variety of experimentally calibrated geobarometers show that metamorphism, and the intrusion of the Fiambalfi gabbro-norite, occurred at 7-8 kbar (20-25 km depth). The distribution of cumulate ultramafic rocks predominantly along the west side of the Fiambalfi gabbro-norite (Fig. 2) suggests that rocks on the west side of the southern block were structurally deeper than rocks to the east during intrusion (DeBari 1994). Pressure estimates generally support this interpretation, although the 1-3 kbar pressure variations are within the accuracy limits of the barometers. Timing of events. The timing of Late Cambrian to Ordovician events in the southern block is constrained by U-Pb geochronological data and 4~ ages on amphiboles. The data, presented and discussed in detail by Grissom (1991, chapter 4), are summarized in Fig. 9. The U-Pb zircon crystallization age of the Fiambal/t gabbro-norite is constrained to 510-515Ma (unpublished data, Table 2). The sample (Z16) was collected from a comagmatic tonalite pod within the Fiambal/t gabbro-norite. A coarse zircon fraction from this sample is nearly concordant at 515 Ma while a fine zircon fraction is nearly concordant at 510 Ma. These ages are within the uncertainty of a Sm-Nd mineral isochron (clinopyroxene, plagioclase, wholerock) age of 5 0 1 + 2 1 M a for a two-pyroxene gabbro (DeBari & Wright unpublished data).

The thermal aureole associated with the Fiambalfi gabbro-norite indicates that emplacement was synchronous with peak metamorphism, and the alignment of granulite-facies minerals to define Sd and La in the vicinity of the Fiambalfi gabbro-norite indicate that deformation continued during intrusion. Hence the U-Pb zircon and Sm-Nd ages of the gabbro-norite date peak contact metamorphism and active deformation at 510-515 Ma. The distinction between regional and contact metamorphism is important in evaluating the timing of Ordovician events in the southern block. The timing of peak contact metamorphism near the Fiambalfi gabbro-norite is not necessarily the same as the timing of peak regional metamorphism. The timing of metamorphism is best evaluated using U-Pb monazite ages as described below for samples Z16, Z6, and Z15 (Table 1, Fig. 9). Monazite from sample Z16 (the tonalite pod in gabbro-norite) yields three different concordant ages for different grain-size fractions (Table 1); the coarsest monazite yields an age of 505 Ma, medium-sized monazite yields an age of 496Ma, and the finest monazite yields an age of 493Ma. All the ages are interpreted as cooling ages; their spread probably reflects different effective closure temperatures to Pb diffusion for different grain sizes. Samples Z6 and Z15 were collected from the migmatitic rocks west of the Fiambal/t gabbronorite. These migmatites were interpreted as in-situ partial melts, formed at temperatures of 800-860~ by biotite dehydration melting during intrusion of the gabbro-norite (Grissom 1991, chapter 3). The monazites from samples Z6 and Z15 yield concordant ages of 473 and 471 Ma respectively, with no significant variation in age related to size fraction (Table 1). These are interpreted as cooling ages through monazite closure to Pb diffusion. The difference of 20-30 Ma between the monazite cooling ages of samples Z6 and Z15 and those for Z16 probably reflects grain size and/or compositional control on the closure temperature. The typical range of monazite closure temperatures is 600-750~ (e.g., Mezger 1990). Conductive cooling requires that any point within the Fiambal/t gabbro-norite had to cool through a given temperature at the same time or later than any point in the country rocks. We therefore interpret the monazite ages as reflecting cooling through 600-650~ for samples Z6 and Z15 and cooling through 650-750~ for sample Z16. The total spread of monazite ages between 505 and 471 Ma suggests relatively slow cooling,

GEOLOGY OF THE SIERRA DE FIAMBALA on the order of 6~ -I, through the range of monazite closure temperatures. This is in contrast to the fast cooling, on the order of 50-100~ -1, from c. 1200~ during gabbronorite intrusion at 510 515Ma to the temperatures of monazite closure in sample Z16 at 493-505Ma. Since regional metamorphic temperatures reached 700~ (Fig. 11), at or above the closure temperature of some monazite grains, the younger monazite ages (c. 470 Ma) probably reflect continued regional metamorphism at T > 650~ for 20-30Ma after the intrusion of the Fiambalfi gabbro-norite. Distributed penetrative deformation also continued after the emplacement of the Fiambal/t gabbro-norite as indicated by the folding and foliation development within the gabbro-norite. The cessation of distributed deformation is constrained by 4~ ages on amphibole. All the amphibole ages from the southern block (Fig. 9) are for amphiboles separated from boudinaged mafic intrusions, except sample 687. The amphibole ages from the mafic intrusions indicate cooling through amphibole closure between 430 and 455 Ma. Sample 687 is a calc-silicate gneiss collected adjacent to the gabbro-norite contact. The peak metamorphic assemblage in this sample includes plagioclase, calcite, quartz, diopside, sphene, and scapolite, but not amphibole. The amphibole occurs as a retrograde phase rimming diopside, with textures indicative of static mineral growth. The step-heating spectrum for this sample defines a plateau with an apparent age of 443 • 1 Ma and an isochron age of 4 4 6 ~ 1 M a . We therefore interpret 443 Ma as a minimum age for the growth of the static retrograde amphibole and the cessation of distributed deformation in the southern block. The structural, metamorphic and geochronological data presented above document parallel histories for the northern and southern blocks during the Cambrian and Ordovician. Greenschist-facies metamorphism and deformation formed a penetrative foliation in the northern block at 540-550Ma, while rocks affected by migmatization reached amphibolite facies. The analogous event in the southern block involved the emplacement of the La Puntilla orthogneiss at e. 550 Ma, accompanied by amphibolite facies metamorphism and the formation of metamorphic layering. Between 550 and 400Ma the northern block was again metamorphosed to greenschist to amphibolite facies (<8 kbar) during folding and the formation of the dominant foliation. The timing of Ordovician events is more tightly constrained in the southern block where the 510-515Ma Fiambal/t gabbro-norite formed a

313

granulite contact aureole during regional amphibolite-facies metamorphism (7-8kbar) that peaked between 505 and 470 Ma. The age of the dominant foliation in the southern block is bracketed between 550 and 440 Ma.

Mid-Palaeozoic

Cooling and retrogression Cooling occurred in both the northern and southern blocks from Ordovician to Devonian times. The southern block experienced localized deformation synchronous with cooling, but no deformational fabrics that overprint Sa were observed in the northern block. The geological and geochronological constraints indicate that structural movements between the northern and southern blocks originated during mid-Palaeozoic cooling. The structural observations associated with localized deformation during cooling in the southern block are summarized in Fig. 12. Calcsilicate rocks southeast of the Ratones granite near Cerro Negro (Fig. 12) have a penetrative foliation that is formed by the alignment of retrograde minerals and flattened quartz grains. The foliation associated with retrogression (Sr) is defined by the alignment of amphibole which grew after diopside and chlorite (amphibolite to greenschist facies). The alignment of amphibole and stretched quartz grains also define a shallowdipping stretching lineation (Lr) that is visible on Sr surfaces. Sr is consistently E W-trending and N-dipping within the structural domain near Cerro Negro (Fig. 12). Farther east along the edge of the range, Sr is rotated in Tertiary fault blocks, and its intensity dies out to W and S. Westwards, along strike with Sr, syn-deformational retrograde textures analogous to those which define Sr are present, but are less pervasive in development. Southwards, these textures die out rapidly within 2-3 km of Cerro Negro. Localized deformation synchronous with retrogression also occurred southwest of the Fiambalfi gabbro-norite. Figure 12 shows a band of mylonitic rocks which occurs at a contact between semipelitic and calc-silicate rocks. The mylonitic shear zone is NW-trending and steeply dipping parallel to the older foliation and lithological contacts. Offset is inferred to be relatively minor across the mylonitic zone since the zone does not disrupt stratigraphy or metamorphic grade. The timing of deformation during cooling is constrained by a 4~ cooling age

314

G. C. GRISSOM ET AL.

Fig. 12. Summary of the structural and metamorphic data related to the mid-to-late Palaeozoic events in the southern block. The grey lines mark the structural domains discussed in the text. Map legend is the same as in Fig. 2 unless otherwise noted. The stereonets are equal-angle lower-hemisphere projections. on amphibole from sample 687 (Fig. 9). The amphibole in this sample has the same optical characteristics and was formed by the same retrograde reaction as the foliated and lineated retrograde amphibole grains near Cerro Negro. Hence, deformation during retrogression was active by 443 Ma but was localized. The minimum age of deformation during cooling near Cerro Negro is approximately 335 Ma based on the age of the Ratones granite (see below), which truncates St.

Formation of the structural break between the northern and southern blocks Several arguments indicate that the discontinuity between the northern and southern blocks is a structural boundary that originated in the mid-late Palaeozoic. The boundary truncates

the monoclinal sequence of semipelitic and calcsilicate rocks in the southern block (Fig. 2) and forms a metamorphic discontinuity between the upper amphibolite facies rocks of the southern block and the greenschist facies schists of the northern block. Movement between the two blocks must have predated intrusion of the Carboniferous Ratones granite since the undeformed granite intrudes the metamorphic discontinuity without offset. The original structure must post-date the 470-500Ma metamorphic assemblages, which are offset, and movement between the northern and southern blocks is thus bracketed between 500 and approximately 330Ma. The southern block gives consistently older 4~ cooling ages on amphiboles (430-455 Ma, Fig. 9) than the northern block (c. 400 Ma), suggesting initial uplift and cooling of the southern block before 430 Ma. However, the 400 Ma ages from amphibole in the northern

GEOLOGY OF THE SIERRA DE FIAMBALA block are minimum ages since these amphiboles may have been disturbed during Carboniferous heating (see below). The timing constraints and the development of Sr near the boundary between the northern and southern blocks suggest that the retrograde foliation and lineation in the vicinity of Cerro Negro may be related to the OrdovicianSilurian displacement between the blocks. However, the structural discontinuity cannot be further characterized because the emplacement of the Ratones granite, the imbrications of the range-bounding fault east of Cerro Negro, and the fault northwest of the Ratones granite have dismembered the older structure. One possible interpretation is that a NW-dipping shear zone juxtaposes rocks of the southern block against quartzo-feldspathic rocks of the northern block. The orientation of the shear zone (based on the orientation of St) would imply a normal sense of displacement. The shear zone is not exposed at the surface due to displacements along Tertiary reverse faults.

Carboniferous Emplacement of post-tectonic granites occurred in the northern and southern blocks during the Carboniferous. Retrogression occurred in metamorphic and hydrothermal aureoles around the plutons. Late, minor folding may have also accompanied the intrusions. The geological and geochronological data associated with Carboniferous events is summarized in Figs 12 and 13. The final folding event in the Sierra de Fiambalfi is characterized by open, wavy folding of Sd and Sr without an associated foliation or cleavage. Open folds with wavelengths of 10-50m and amplitudes of 1-10m were recognized in the northern and southern blocks but are best developed and best documented in a structural domain northeast of the Fiambalfi gabbro-norite (Fig. 12). The measurements of outcrop-scale folds with variably orientated and steeply plunging axes are shown on a stereonet plot in Fig. 12. Within this domain Sd is folded on a kilometre-scale around the southeastern lobe of the Ratones granite into an E-W-trending orientation (Fig. 8b). The lower stereonet plot in Fig. 12 shows that the poles to Sd define a rr-axis within the spread of measured fold hinges. Rare open folds of Sr in the vicinity of Cerro Negro were also observed, indicating that the late folds post-date Sr. Notice that the poles to Sr plotted in the upper stereonet define a re-axis close to the cluster of fold hinges shown in the lower stereonet plot, although the two

315

plots are for different domains. The open folds are interpreted as post-dating S r and may be related to the emplacement of the post-tectonic plutons, although the lack of mineral growth associated with the folds makes this hypothesis difficult to evaluate. The final thermal event in the Sierra de Fiambalfi is associated with the emplacement of post-tectonic granites and related aplitic or pegmatitic dykes. The undeformed plutons and dykes intrude all lithologies in the northern and southern blocks with relatively sharp intrusive contacts that cut all deformational fabrics. Retrograde mineral growth and hydrothermal alteration with static textures is spatially associated with the intrusive bodies. Retrograde mineral growth in the southwest part of the range includes randomly orientated fine-grained shredded chlorite after mafic minerals and sericitization of feldspars in all lithologies. The sericite and chlorite are usually associated with quartz and calcite veins that cut all foliations. In the southwest part of the range, late randomly-orientated muscovite only occurs within 0.5 km of post-tectonic granite contacts or near swarms of undeformed aplitic and pegmatitic dykes shown in Fig. 12. To the northeast, retrograde muscovite is more common and retrogression is generally stronger. In the eastern part of the northern block quartz grains are completely annealed whereas sub-grains are preserved in quartz to the south and west. The post-tectonic plutons include biotite granite and biotite hornblende granite that are rich in microcline and occasionally contain muscovite. The largest intrusion is the Ratones granite, which intrudes the contact between the northern and southern blocks (Fig. 2). This is a composite body containing an equigranular biotite granite phase, transitional to a porphyritic biotite+amphibole granite phase with microcline phenocrysts (Arrospide 1984). Small bodies of equigranular biotite granite intrude the Fiambalfi gabbro-norite (Fig. 2) and two larger bodies of porphyritic to equigranular biotite 4- amphibole granite occur further to the southwest. Pegmatitic, aplitic and rhyolitic dykes both emanate from and crosscut the granites. 4~ and U - P b geochronological data constrain the timing of post-tectonic intrusions and related retrograde mineral growth. 4~ data are presented in Appendix B and U-Pb data in Table 1. Sample locations and step-heating spectra are shown in Fig. 13. U - P b ages on monazite and zircon separated from samples of post-tectonic granites constrain the crystallization ages of the intrusion. Sample

316

G. C. GRISSOM E T A L .

Fig. 13. Sample locations and summary of U-Pb and 4~ geochronological data related to post-tectonic intrusions and the Late Palaeozoic reheating event. The map is a simplified version of Fig. 2. Z4 is an equigranular biotite granite from the southwest margin of the Ratones granite (Fig. 13). U-Pb analyses of three fractions give discordant ages that indicate a minimum intrusive age of 335Ma (the lowest 2~176 age). Sample Z3 is a porphyritic biotite hornblende granite from the pluton in the south central portion of Fig. 13. Monazite from this sample gives a reasonably concordant age of c. 325Ma, which may represent minor cooling after crystallization. 4~ ages from step-heating experiments on mineral separates (two muscovites, three biotites and a microcline) are presented below and summarized in Fig. 13. They are interpreted as cooling ages or reheating ages, since Ordovician metamorphic temperatures exceeded the temperatures of retention or blocking for argon diffusion in these minerals. Muscovite separated from an amphibolitefacies schist in the southern block (sample 330) and a pegmatitic dyke within the migmatized schist of the northern block (sample 651) yield consistent ages. Both have muscovite and biotite aligned with Sd, and later penetrative foliations are absent. The Ar age spectrum for sample 330 has no plateau, due to minor excess 4~ but the isochron gives a well-defined apparent age of

368 + 1 Ma. Sample 651 shows a discordant stepheating spectrum that suggests the presence of excess 4~ together with superposed Ar-loss; the isochron age is well constrained at 360 + 2 Ma. The minimum age of cooling through c. 300~ is taken as 369 + 1 Ma for sample 330 and 360 Ma for sample 651. In both samples, the 320-330 lower intercepts are interpreted as reflecting partial degassing during the emplacement of post-tectonic granites. Biotite was collected from the southwestern margin of the La Puntilla orthogneiss (sample Z1), a biotite feldspar gneiss from the northern block (sample Z12), and the Ratones granite (sample Z4) (Fig. 13). For sample Z1, the step heating spectrum shows a well defined plateau at 355 + 1 Ma with a decrease in ages trending towards 300-320 Ma at low temperatures. 355 Ma is interpreted as the minimum age for cooling through c. 280~ The spectrum for Z12 forms a plateau at 332.0+0.5Ma with minor Ar loss. The lower intercept for sample Z1 and the 332Ma plateau age for sample Z12 are thought to reflect resetting of the Ar isotope system during the emplacement of the post-tectonic granites. Sample Z4, from the Ratones granite, does not give a plateau age, but an isochron plot yelds an apparent age of

GEOLOGY OF THE SIERRA DE FIAMBALA 336 i 3 Ma, consistent with the U-Pb minimum age of 335 Ma for the same sample. The step-heating spectrum for microcline in sample Z1 (described above) is also shown in Fig. 13 and shows a smooth increase from about 300 Ma to 330 Ma. As with the muscovite and biotite date, the spectrum is interpreted as reflecting complete cooling after the intrusion of post-tectonic granites. The complete resetting of the orthogneiss microcline in sample Z1 indicates that temperatures in the southwest portion of the range during the Carboniferous reheating event exceeded c. 150~ the minimum estimated closure temperature for microcline (Harrison & McDougall 1981). The 4~ and U-Pb geochronological data confirm that static retrograde mineral growth and veining was related to undeformed intrusions, since ages were either reset to, or give lower intercepts near, the crystallization age of the post-tectonic granites. The 325 Ma monazite age on sample Z3, the resetting of the biotite age in sample Z12 to 332Ma, and the c. 330Ma biotite age for sample Z4, all suggest that the intrusive age of the Ratones granite is not significantly older than the 335Ma minimum age determined on zircon from sample Z4. The muscovite and biotite ages older than this require that cooling below approximately 300~ in the northern and southern blocks occurred prior to granite emplacement. The evidence for argon loss in all spectra requires that plateau ages be interpreted as minimum cooling ages and precludes a comparison of the cooling histories for the two blocks. As inferred from the observations of retrograde mineral growth, Carboniferous reheating and retrogression were more pervasive toward the northeast part of the range.

Discussion Field mapping in conjunction with structural, metamorphic, and geochronological data from the Sierra de Fiambalfi identify three major tectono-thermal events and three important structural-stratigraphical units within the Sierra de Fiambalfi. The three tectono-thermal events include: (1) intrusion, migmatization, deformation, and metamorphism in the Early Cambrian, (2) mafic intrusion, deformation, and metamorphism in the Late Cambrian to Ordovician, and (3) reheating, retrogression, and posttectonic granite emplacement in the Carboniferous. In addition, the preliminary U-Pb data on sample Z11 indicate that the gneisses and migmatites at the eastern side of the northern block

317

may have a Precambrian intrusive-metamorphic history. The three structural-stratigraphical units are: (1) the Precambrian (?) gneisses and migmatites on the east side of the northern block, (2) the Palaeozoic schists at the western side of the northern block, and (3) the metasedimentary rocks of the southern block. The present data show that all three structural-stratigraphical units were affected by the same Cambrian, Ordovician, and Carboniferous tectono-thermal events. Several arguments suggest that the discontinuity between the gneiss and schist in the northern block is a structural or depositional contact between Precambrian continental crust (South American craton) and Late Precambrian to Palaeozoic sedimentary rocks. If orthogneiss sample Z l l has a Precambrian crystallization age, it infers a Precambrian metamorphic and deformational history for the adjacent paragneiss at the east side of the northern block. The gneissic layering that predates Early Cambrian migmatization (Fig. 3) may be Precambrian in age. In contrast, only Early Cambrian and Ordovician structures and thermal events were recognized in the schist on the west side of the northern block and in the metasedimentary rocks of the southern block. Hence, schist of the northern block may be cover-rock deposited over Precambrian craton represented by the gneiss of the northern block. This hypothesis has important implications for the regional geology; however, more detailed mapping of the schist-gneiss contact and the completion of U-Pb analyses on sample Z ll will be necessary in order to further evaluate the idea. The discontinuity between the northern and southern blocks juxtaposes upper amphibolitegranulite facies rocks of the southern block and greenschist to amphibolite-facies rocks of the northern block. The contact between the northern and southern blocks also marks a stratigraphical discontinuity between interlayered semipelitic and calc-silicate rocks in the southern block and the quartzo-feldspathic schists of the northern block. The c. 400 Ma 4~ amphibole cooling ages from the northern block in contrast to the 430-450 Ma 4~ amphibole cooling ages from the southern block suggest that a Late Ordovician structure accommodated relative uplift of the southern block. This structure is tentatively interpreted as a shear zone with a normal sense of motion. However, this interpretation is highly speculative due to the subsequent emplacement of the Ratones granite along the contact, and the complications caused by Tertiary faults that cut across the older structure.

318

G. C. GRISSOM E T A L .

Regional geological implications Sedimentary and volcanic sequences from the Palaeozoic continental margin of South America (Gondwana) occur in the Puna plateau, the fold and thrust belts inboard of the Puna plateau, and in the Famatina System (Fig. 1). The Neogene reverse fault uplifts in the northern Sierras Pampeanas, including the Sierra de Fiambalfi, expose medium- and high-grade metamorphic rocks inferred to be deeper structural levels of the same continental margin (Willner & Miller 1986; Willner et al. 1987). Grissom et al. (1992) argue that the tectonothermal history of the Sierra de FiambalA is correlative with the events documented in other ranges of the northern Sierras Pampeanas. In addition, the emplacement of the Fiambalfi gabbro-norite may be correlated with Ordovician arc volcanism in adjacent exposures of Palaeozoic sedimentary and volcanic rocks (see also Bahlburg this volume and Saavedra et al. this volume). The geochronological data presented here allow a more systematic comparison of the Cambrian and Ordovician events in the Sierra de Fiambalfi and in other parts of the northern Sierras Pampeanas. Similarly, the timing of events in the exhumed deep crustal rocks of the northern Sierras Pampeanas can be compared with sedimentary, volcanic, and lowgrade metamorphic events recognized from the supracrustal sequences in the adjacent tectonic provinces. Figure 14 summarizes the Palaeozoic history of the Sierra de Fiambalfi, the northern Sierras Pampeanas, and the adjacent tectonic provinces containing supracrustal rocks. The column for the northern Sierras Pampeanas in Fig. 14 represents data compiled from the literature for the Sierra de Ancasti and the Sierra de Quilmes (Fig. 1). A more extensive description of the tectono-thermal events documented in these ranges is provided in Grissom et al. (1992) and in the references for Fig. 14. The column shows that the timing of tectono-thermal events is comparable in the Sierra de Fiambal/t, the Sierra de Ancasti, and the Sierra de Quilmes. More detailed study of these ranges will undoubtedly illustrate local variations in the character and timing of regional events. However, the generally correlative histories for ranges in the northern Sierras Pampeanas confirms that these rocks record tectonic events of regional extent. The generalized stratigraphical columns in Fig. 14 for the Famatina System, the Puna plateau, and the fold and thrust belt inboard of the Puna allow a systematic comparison with the tectono-thermal events in the northern Sierras Pampeanas. The Puncoviscana Formation,

which forms the basal unit of the supracrustal exposures (Fig. 14), consists of psammites and local carbonates interpreted as a late Precambrian to Cambrian passive margin sequence (Baldis & Omarini 1984; Je~ek & Miller 1987). The base of the Puncoviscana Formation has not been observed, although Je~ek et al. (1985) argue, on the basis of provenance data, that the psammites rest on extended Precambrian crust of the Brazilian craton. Lithological, chemical, and palaeogeographical correlations between the Puncoviscana Formation and the metasedimentary rocks of the Sierras Pampeanas (Acefiolaza & Toselli 1981; Je~ek & Miller 1987; Je~ek et al. 1985; Willner et al. 1985) have been used to infer that the rocks in the Sierras Pampeanas are structurally deeper equivalents of rocks in the Puncoviscana Formation. During the Cambrian the Puncoviscana Formation was affected by metamorphism, deformation, and intrusion. Early Cambrian metamorphism and deformation in the Sierra de Fiambal~t and the northern Sierras Pampeanas was contemporaneous with sub-greenschist- to greenschist-facies metamorphism in the Puncoviscana Formation dated at 525-570 Ma (K-Ar, whole-rock, Acefiolaza & ToseUi 1981). The 519-534 Ma Cafiani and the 536 + 7 Ma Tastile plutons (Bachmann et al. 1987) in the Puna and fold and thrust provinces (Fig. 1) cut the regional foliation in the Puncoviscana Formation and form aureoles that overprint the low-grade regional metamorphism. These calcalkaline, mildly peraluminous plutons emplaced in the supra-crustal exposures are correlative with the mildly peraluminous La Puntilla orthogneiss and the other 540-550Ma orthogneiss bodies in the Sierra de Fiambalfi. These plutons may be related to the Early-Mid-Cambrian collision between continental terranes (Pampean orogeny) described by Rapela et al. (this volume). In the Late Cambrian to Ordovician the rocks in the Sierra de Fiambalfi experienced a second metamorphic event accompanied by the intrusion of subduction-related magmas (i.e., Fiambalfi gabbro-norite). These subduction-related magmas are equivalent to the Famatinian subduction event that started c. 500 Ma and probably lasted until 455Ma (Pankhurst et al. this volume). A contemporaneous thermal and magmatic history is recognized in the rest of the northern Sierras Pampeanas (Fig. 14). The correlation between Ordovician magmatism and metamorphism in northern Sierras Pampeanas and the contemporaneous sedimentary-volcanic sequences in the Famatina System, the Puna plateau, and the fold and thrust belt inboard of the Puna are discussed in Grissom et al. (1992).

GEOLOGY OF THE SIERRA DE FIAMBALA Northern Sierras Pampeanas

Sierra de Fiambahi

Famatina

Puna

System

Plateau

319

Fold & Thrust Belt

Siltstones and greywackes Marine shales

Carbon-Ma-iferous

I lll[l[l[lI[

Sandstones Carbonates and marls

Devonian :i

Ma--

"

Silurian

- , ' : - - : - : "5. . . . . !

:-": _ _ _ _-_ _ _ _ _-_ -. -.-

I

~]

Conglomerates

~

Hiatus and unconformity

0

Granitic intrusion s

~ 7

with radiometric ages

~Jplittof SouthernBloct >440

Ordovieia l

Mafic intrusions

,

"

0

~

Ophiolitic fragments

500 M a ~ 5~15 metamorphism with abbreviated grade and age

Cambrian 540-550 Ma

Late PC

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Fig. 14. Schematic stratigraphical columns showing the correlative histories for the Sierra de Fiambalfi and the major tectonic provinces of northwestern Argentina. The locations of the tectonic provinces are shown in Fig. 1. The ellipsoidal shapes of intrusive bodies represent the time-span of intrusive activity as constrained by radiometric ages and field relations. Compiled from Turner (1970), Schwab (1973), Furque & Cuerda (1979), Turner & Mendez (1979), and Acefiolaza & Toselli (1984).

The observations in the Sierra de Fiambalfi suggest that the southern block was uplifted in the Late Ordovician to Silurian, between 440 and 470 Ma, the same time as the Ocloyic phase of deformation (Turner & M6ndez 1979) recognized in the Puna Plateau (Fig. 14). Ocloyic deformation in the Puna resulted in folding of the Ordovician sediments and sub-greenschist facies metamorphism (Schwab 1973). The deeper levels of exposure in the northern Sierras Pampeanas relative to the adjacent tectonic provinces suggests that Late Ordovician uplift and exhumation may have been more extensive in the northern Sierras Pampeanas than in the adjacent provinces. For instance, upper amphibolite facies rocks are juxtaposed with greenschist facies rocks at the boundary between the northern and southern blocks in the Sierra de Fiambals suggesting a significant amount of displacement between the two blocks. However, Devonian-Carboniferous mylonite zones in other ranges of the Sierras Pampeanas (Caminos 1979; Miller & Willner 1981) and the deposition of thick conglomerate sequences in the adjacent provinces (Fig. 14) suggest an additional period of uplift in the late Palaeozoic. In any case, the

deep crustal rocks of the northern Sierras Pampeanas were at near surface conditions by the late Palaeozoic (Fig. 14), as indicated by apatite fission track ages (Jordan et al. 1989), the formation of an erosional peneplain surface often covered by Carboniferous sediments (Gonzlez Bonorino 1950), and the 4~ cooling ages on microcline, biotite, and muscovite from the Sierra de Fiambalfi. The correlative timing and character of early Palaeozoic events in the northern Sierras Pampeanas, the Puna, the fold and thrust belt, and the Famatina System (Fig. 1) support the interpretation that the metamorphic and intrusive rocks of the northern Sierras Pampeanas represent deep structural levels of the early Palaeozoic Gondwana continental margin. However, a tectonic explanation of this uplift awaits a better characterization of mid-to-late Palaeozoic structural events in the northern Sierras Pampeanas. Conclusions

All rocks in the Sierra de Fiambal~t experienced two periods of metamorphism, deformation, and

320

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GEOLOGY OF THE SIERRA DE FIAMBALA intrusion; first during the Cambrian (Pampean orogeny) and again during the Ordovician (Famatinian orogeny). The Cambrian event involved granitic intrusion at 540-550 Ma associated with lit-par-lit injection and amphibolitefacies metamorphism in the migmatites of the northern block, while greenschist-facies metamorphism accompanied the formation of a penetrative foliation (SD in the schist to the west. In the southern block the c. 550Ma intrusion of the La Puntilla orthogneiss was accompanied by amphibolite-facies metamorphism and the formation of compositional layering in the metasedimentary rocks. Ordovician events in the northern block involved mafic intrusion, a second metamorphic event, and the formation of the dominant foliation (Sd) between 550Ma and c. 400Ma. The correlative event in the southern block involved the intrusion of the Fiambal~ gabbronorite at c. 500-515Ma, followed by amphibolite to granulite-facies metamorphism between 515 and 470 Ma, with contemporaneous formation of the dominant foliation (Sd). The rocks in the southern block were probably uplifted relative to northern block in the Late Ordovician and all locations experienced cooling through mid-late Palaeozoic times. Carboniferous reheating and retrograde metamorphism associated with the emplacement of post-tectonic granites at 325 to c. 340Ma was most pervasive toward the northeast part of the range. The Cambrian, Ordovician, and Carboniferous events in the Sierra de Fiambal~ were of regional extent as indicated by temporal correlation with other parts of the northern Sierras Pampeanas. Correlation between regional events in the northern Sierras Pampeanas and events in the adjacent supracrustal exposures confirms that rocks in the northern Sierras Pampeanas formed at deep (10-25km) structural levels in the early Palaeozoic continental margin of Gondwana. We acknowledge the support of J. E. Wright, who provided the expertise and facilities for the U-Pb geochronological work. L. Villar initiated the Sierra de Fiambalfi project and provided logistical support through the Direcci6n Nacional de Geologia y Mineria, Argentina. This work would not have been possible without the field support of S. Araya, A. Araya and F. Saavedra whose knowledge of mules, mountains and water made field work in a dry and desolate place enjoyable. Financial support was provided by a National Science Foundation Grant INT#8613115 to R. G. Coleman. The manuscript benefited greatly from reviews by I. U Millar, R. J. Pankhurst and an anonymous reviewer.

321

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JOHNSON, M. C. & RUTHERFORD, M. J. 1989. Experimental calibration of the aluminium-inhornblende geobarometer with application to Long Valley caldera (California) volcanic rocks. Geology, 17, 837-841. JORDAN, T. E., ZEITLER, P., ALLMENDINGER, R. W., RAMOS, V. A. & GLEADOW,A. J. W. 1989. Thermochronometric data on the development of the basement peneplain in the Sierras Pampeanas, Argentina. Journal of South American Earth Sciences, 2, 207-222. KNOVER, M. & MILLER, H. 1982. Rb-Sr geochronology of the Sierra de Ancasti (Pampean Ranges, NW Argentina), 5 ~ Congreso Latinoamericano de Geologia, Buenos Aires, Actas, 457-471. KROGH, T. E. 1973. A low-contamination method for hydrothermal decomposition of zircon and extraction of U and Pb for isotopic age determinations. Geochimica et Cosmochimica Acta, 46, 637-649. LOVERA, O. M., RICHTER, F. M. & HARRISON, T. M. 1989. 4~ thermochronometry for slowly cooled samples having a distribution of diffusion domain sizes. Journal of Geophysical Research, 94, 17917-17936. MERRIHUE, C. & TURNER, G. 1966. Potassium-argon dating by activation with fast neutrons. Journal of Geophysical Research, 71, 2852-2857. MEZGER, K. 1990. Geochronology in granulites. In: VIELZEUF, D. & VIDAL, PH. (eds) Granulites and Crustal Evolution. Kluwer Academic Publishers, Netherlands, 451 470. MILLER, H. • W1LLNER, A. D. 1981. The Sierra de Ancasti (Catamarca Province, Argentina) an example of polyphase deformation of Lower Paleozoic age in the Pampean Ranges. Zentralblatt ffir Geologie und Palontogie, Teil I, (3/4), 272-284. NEWTON, R. C., CHARLU, T. V. & KLEPPA, 0. J. 1980. Thermochemistry of high structural state plagioclase. Geochimica et Cosmochimica Acta, 44, 933-941. PANKHURST, R. J., RAPELA, C. W., SAAVEDRA, J., BALDO, E., DAHLQUIST, J., PASCUA, I. & FANNING, C. M. 1998. The Famatinian magmatic arc in the central Sierras Pampeanas. This volume. PENCK, W. 1920. Der Siidrand der Puna de Atacama (NW-Argentinien). Ein Beitrag zur Kenntnis des andinen Gebirgstypus und zu der Frage der Gebirgsbildung.). Abhandlungen der Sdchsischen Akademie der Wissenschaften, Leipzig, mathematisch physikalische Klasse, 37, 1-33. RAPELA, C. W. 1976. E1 basamento metam6rfico de la regi6n de Cafayate, provincia de Salta. Aspectos petrol6gicos y geoquimicos. Revista de la Asociaci6n Geol6gica Argentina, 31, 203-222. - - , HEAMAN, L. M. & MCNUTT, R. H. 1982. Rb-Sr geochronology of granitoid rocks from the Pampean Ranges, Argentina. Journal of Geology, 90, 574-582. --, PANKHURST, R. J., CASQUET, C., BALDO, E., SAAVEDRA, J., GALINDO, C. & FANNING, C. M. 1998. The Pampean Orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de C6rdoba. This volume.

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-

The granitoids of the Sierra de San Luis E. J. L L A M B I A S 1, A. M. S A T O 1, A. O R T I Z

S U , ~ R E Z 2 & C. P R O Z Z I 3

1 Universidad Nacional de La Plata, Centro de Investigaciones Geoldgicas, Calle 1 No. 644, 1900 La Plata, Argentina 2 Universidad Nacional de San Luis, Departamento de Geologia, Chacabuco y Pedernera, 5700 San Luis, Argentina 3 Universidad Nacional del Sur, Departamento de Geologia, San Juan 670, 8000 Bahia Blanca, Argentina Abstract: The Sierra de San Luis consists of Lower Palaeozoic igneous and metamorphic

rocks. The granitoids were emplaced between Late Cambrian and Early Carboniferous times, and the most important metamorphic and deformational events occurred during this period. Older metamorphic events have also been recognized but they are poorly constrained. The most important group of granitoids are as follows: (1) Deformed tonalites and granites. The tonalites have a calc-alkaline signature with low La/Yb ratio. They are the Early Famatinian, pre-orogenic, granitoids of Late Cambrian to Early Ordovician age and form part of a magmatic arc. (2) The La Escalerilla granite is a deformed biotite monzogranite and consists of a single elongated pluton; its relation with the other granitoids is not clear. (3) Garnet-biotite-muscovite granodiorites and granites and pegmatites. The granitoids have S-type signatures, with high 87Sr/86Sr and La/Yb ratios. They are possible crustal magmas related to a thickened crust and are interpreted as early- to syn-orogenic. (4) Biotite- and biotite-hornblende granites and minor tonalites of Devonian to Early Carboniferous age. The granites form large, post-orogenic, plutons and batholiths related to crustal uplift, and represent the last igneous activity of the Famatinian cycle (and also of the Palaeozoic era) in the Sierra de San Luis.

The Sierra de San Luis is part of the Sierras Pampeanas tectono-stratigraphical province, which includes the Upper Proterozoic and Lower Palaeozoic metamorphic and igneous rocks of central Argentina. The present topography is the result of Neogene tectonics (GonzMez Bonorino 1950; Jordan & Allmendinger 1986). Based on the distribution of metamorphic and igneous rocks, the Sierras Pampeanas province was divided by Caminos (1972, 1979) into the Western and Eastern Sierras Pampeanas. The former are characterized by abundant mafic, ultramafic and calcareous rocks, whilst in the latter mafic and ultramafic rocks are found only in minor belts, and tonalitic to granitic rocks are more abundant. The first regional characterization of metamorphism and structures of the Western and Eastern Sierras Pampeanas was presented by Dalla Salda (1987). Isotopic dates ranging from 0.9 to 1.1 Ga for the Western Sierras Pampeanas (see synthesis in Varela et al. 1996) supported the division made by Caminos (1972, 1979). The Sierra de San Luis and the Sierras de Cdrdoba make up the southern extreme of the Eastern Sierras Pampeanas. In the synthesis of Gordillo & Lencinas (1979) these two mountain blocks were described separately because of the higher metamorphic grade in the Sierras de

C6rdoba. The ages published by Cingolani & Varela (1975) provided evidence for a Late Precambrian tectono-thermal event in the Sierras de C6rdoba, although new isotopic ages point to a Cambrian metamorphic event (Rapela et al. and Sims et al. this volume). In the Sierra de San Luis this older orogeny is not well known and is overprinted by Famatinian orogenic phases (Llambias et al. 1996b). Despite the differences in lithology and age of igneous and metamorphic events within individual mountain blocks, a distinctive feature of the entire Sierras Pampeanas province is its exhumation prior to the deposition of Upper Carboniferous to Permian continental sediments, collectively named the Paganzo Group. Granitic rocks emplaced at different times are widely distributed in the Sierra de San Luis. Llambias et al. (1991) and Ortiz Su~irez et al. (1992) defined them as pre-, syn-, and postkinematic with respect to the main deformational event. In this paper the granitoids are classified on the basis of the shape of the intrusive bodies, their composition, and their relationships to the deformation and metamorphism of the country rocks (see von Gosen & Prozzi this volume). The main deformational event was described by Kilmurray & Dalla Salda (1977) and Criado Roqu6 et al. (1981) as that which produced the

LLAMBiAS,E. J., SATO,A. M. ORTIZSUAREZ,A. & PROZZI, C. 1998. The granitoids of the Sierra de San Luis. In: PANKHURST, R. J. & RAPELA,C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 325-341.

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Fig. 1. Simplified geological map of the Sierra de San Luis, modified from Ortiz Su{trez et al. (1992) with the location of the main granitic outcrops.

GRANITOIDS OF THE SIERRA DE SAN LUIS present structures of the metamorphic rocks, and is referred to the Ocloyic phase (Turner & M6ndez 1975, 1979). The age of this event has been assigned to the Ordovician by these authors and forms part of the Famatinian orogenic cycle. This cycle was defined by Acefiolaza & Toselli (1973) as encompassing the geological evolution of the central and northwestern Argentina between Mid-Cambrian and Late DevonianEarly Carboniferous times. The lower limit of the Famatinian cycle is indicated by an angular unconformity between the Mes6n Group (Middle to Upper Cambrian) and the underlying Puncoviscana Formation (Upper ProterozoicLower Cambrian) in NW Argentina. In the Sierra de San Luis this lower limit cannot be observed because of the intensity of later metamorphic and igneous activity. We consider the tonalitic intrusions to be the first manifestations of this orogenic cycle. The country rocks, consisting of both gneisses (Llambias et al. 1996b) and previously undeformed sedimentary rocks (von Gosen 1998, in press), are considered to belong to older orogenic cycles, such as the Pampean cycle (Acefiolaza & Toselli 1973) or the Brasiliano cycle (Almeida 1969, and see discussion in Trompette 1994, p. 111). Prozzi (1990) has pointed out the possible correlation of the above mentioned undeformed sedimentary rocks of San Luis (albeit later affected by Famatinian metamorphism and deformation, together with the granitoids) with the sediments of the Puncoviscana Formation. The deposition of the Upper Carboniferous continental sediments of the Paganzo Group fixes the upper limit of the Famatinian orogenic cycle. Exhumation of the metamorphic and igneous rocks probably occurred around the Devonian-Carboniferous boundary.

S u m m a r y o f the m e t a m o r p h i c rocks

In this paper the metamorphic rocks of the Sierra de San Luis are divided into four units (Fig. 1), according to the scheme of Ortiz Sufirez et al. (1992). For more details and relationship to structural evolution see von Gosen & Prozzi (this volume). Gneisses. We distinguish two types of gneisses. (1) Biotite-garnet gneisses, partially migmatitic, containing quartz, plagioclase (andesine), sillimanite, garnet, K-feldspar. Associated with these gneisses there are lens-like amphibolite bodies and mafic to ultramafic intrusive rocks. Metamorphism reached the highest levels of the amphibolite facies. (2) Muscovite-

327

biotite gneisses, commonly injected by quartzoligoclase veins. The gneisses contain muscovite, biotite, quartz, plagioclase (oligoclase-andesine) and in some cases garnet. The metamorphic grade corresponds to amphibolite facies. Mica-schists. These mainly consist of quartz, muscovite, biotite (garnet, plagioclase), with local growth of andalusite, staurolite and rarely cordierite and sillimanite porphyroblasts. On the eastern side of the Sierra de San Luis the mica-schists contain quartz-oligoclase veins. The metamorphic grade reached lower amphibolite facies. Phyllites. These consist of phyllites, with minor schists and slates. The metamorphic grade corresponds to greenschist facies. The protolith consisted of shales, sandstones, conglomerates, and minor silicic igneous rocks.

Early Famatinian pre-orogenic granitoids: Late Cambrian-Early Ordovician Within this granitoid group we include intrusive bodies that share the deformation of the country rock. The most abundant rocks are of intermediate compositions, mainly tonalites, described as the tonalite group. Similar structural characteristics are found in a few granitic bodies, which are described as the granite group. The tonalite group

Outcrops of tonalitic rocks are abundant in the western part of the Sierra de San Luis, extending from south of the Pancanta valley to E1 Realito in the north. Most of the plutons are distributed along a NE-trending ductile shear belt that separates biotite-muscovite gneisses in the west from phyllites in the east. They are probably the remnants of a large batholith. Only some of the plutons have been studied in detail, e.g., Bemberg, Las Verbenas and Gasparillo (Fig. 2) in the Pancanta Valley (Sato 1993; Llambias et al. 1996b; Sanchez et al. 1996), and E1 Realito at the NW extreme of the range (Brogioni et al. 1994). The Tamboreo pluton (Zardini 1966) is a small, isolated tonalitic pluton located to the east (Fig. 3) along the contact of mica-schists and phyllites. Its structural relationship with the country rocks, its shape, and its composition lead us to include this pluton in the tonalite group. In the western belt, the tonalitic rocks were intruded into already metamorphosed and

328

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tary rocks that were later transformed to phyllites (von Gosen 1998, in press). Contact aureoles with porphyroblasts of biotite, andalusite and muscovite are still recognized, despite

the subsequent deformation and greenschistfacies metamorphism produced by the main orogenic phase of the Famatinian cycle. The studied plutons are generally small, up to 6 km 2 in outcrop, with the major axis <8 km in length and maximum aspect ratios of 3:1. The E1 Realito pluton (Fig. 1) is an exception, reaching a length of 23 km and an aspect ratio of almost 10:1. The original shapes of the plutons have been modified by the deformation, producing highly sheared borders and less deformed interiors. The most abundant rock is a medium-grained tonalite, transitional to diorite and quartz gabbro or to granodiorite and monzogranite (Fig. 4a). Zoned plagioclase, quartz, K-feldspar and biotite are the main minerals, whilst amphibole is more abundant in mafic facies. Mafic microgranular enclaves are common and consist of biotite, amphibole, plagioclase and quartz. The tonalites are also intruded by different pulses of dykes of tonalitic, granitic, aplitic, and pegmatitic compositions. Some of these dykes could be related to the later granitic intrusions, such as the La Escalerilla granite.

GRANITOIDS OF THE SIERRA DE SAN LUIS

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The regional foliation of the country rocks can be traced within the plutons with an irregular distribution, and also affects the mafic enclaves and acidic dykes. Within the tonalites the foliation is defined by the orientation of biotite-muscovite-epidote aggregates, lensoid aggregates of quartz, and lesser amounts of recrystallized feldspars and amphibole. The Tamboreo pluton (Fig. 3) has a pear-like shape (6km 2) and consists mainly of tonalite transitional to granodiorite. It contains well zoned plagioclase, abundant biotite, quartz and scarce K-feldspar, and amphibole. The modal abundance of amphibole decreases with the degree of deformation to yield more biotite, epidote, and quartz. Microgranular mafic enclaves are common. A contact aureole is recognised along the northern and southeastern contacts; it is rich in fine-grained biotite and microporphyroblasts entirely altered to sericite, quartz, and pennite (?), which are interpreted as original andalusite and cordierite. In general, the regional foliation wraps around the intrusive body, although the northern and southeastern contacts are cross-cut by the country rock foliation, which penetrates into the pluton with irregular distribution. A rhyolitic dyke, similar to metavolcanic rocks defined by Brodtkorb et al. (1984) and chemically analysed by Hack et al. (1991), cuts across the western margin of pluton and the mica-schists (Fig. 3). This dyke was deformed together with the tonalite and the country rock (von Gosen 1998). A Rb-Sr whole rock errorchron of 513Ma with large error (MSWD=5.5) was

Age.

reported by Sato et al. (1996) for the Bemberg pluton. Excluding the four most dispersed samples among the 12 analysed, the remaining eight samples delineate an isochron of 512:k16Ma ( M S W D = l . 3 4 ) , with an initial 87Sr/86Sr of 0.7083 + 0.0001 (Sato, unpublished data). Therefore, we consider this last age as a better estimate of the crystallization age. On the other hand, U-Pb SHRIMP dating of zircon from one of the same samples of the Bemberg pluton yielded a main population at 468 ~ 6 M a (Sims et al. this volume), an age younger than that obtained by the Rb-Sr method. At the present level of knowledge, we cannot evaluate this difference. Another similar age of 472 + 5Ma was also obtained for the Tamboreo pluton, by the same method (Sims et al. this volume). C h e m i c a l characteristics. The tonalitic rocks define a subalkaline suite with an SiO 2 range of 48-67%, but mostly 62-67% (Sato et al. 1996). Although modal compositions indicate a clear metaluminous signature, chemical diagrams point to a metaluminous to weakly peraluminous character. This apparent contradiction may be explained by the chemical changes that occurred during deformation and metamorphism. The weak peraluminous tendency is accompanied by a general increase in the K20 content. The trace element signature is typical of calcalkaline series, with a moderate LIL-element enrichment and HFS-element depletion. The normalized REE diagram shows a moderate Eu anomaly and low (7.8-9.7) La/Yb ratios (Fig. 5d). For the tonalite of E1 Realito,

330

E. J. LLAMBIAS E T AL.

Brogioni et al. (1994) mentioned a similar SiO2 range of 56.8-66.0% (tonalites to minor granodiorites), a metaluminous to weakly peraluminous character (A/CNK=0.82-1.15) and trace element abundances typical of magmatic-arc granitoids. The emplacement of dominantly tonalitic plutons with minor gabbros, granodiorites and monzogranites in an elongated belt, and the frequent occurrence of mafic microgranular enclaves, as well as the geochemical signature, are consistent with a subductionrelated magmatic arc. The trend displayed by the analysed samples in the multi-cationic R1-R2 diagram of Batchelor & Bowden (1985) also supports this idea (Fig. 5a). The rather high initial 87Sr/86Sr ratio suggests involvement of crustal rocks in the genesis of the magma.

and garnet as modal minerals. In the R1-R2 diagram (Fig. 5a), the two samples are aligned with the tonalite group. For the monzogranites of E1 Realito, Brogioni et al. (1994) reported a SiO2 range of 70.7-75.0%, a peraluminous character ( A / C N K = 1.11-1.42) and R b - Y - N b relations characteristic of syn-collision granites.

Discussion.

Up to now there is no reported age for these deformed bodies of granitic composition. Although some of them, such as the granite of E1 Realito, intrude the tonalites and are structurally affected by the same deformation, chemically they do not seem to represent a single magmatic evolution. However, more detailed studies are necessary to relate these deformed granitic bodies to the magmatic evolution of the Early Famatinian granitoids.

Discussion.

La Escalerilla pluton The granite group

Only a few granitic bodies share the same structural features of the tonalite group. The Rio Quinto granite is located to the southeast of E1 Tamboreo and comprises a few lens-like sheets less than 1 km 2, parallel to the foliation of the country rocks (Carugno Durfin et al. 1992). The smaller bodies are a few metres thick and 30 m long, and are strongly foliated. The largest body is 800 m long and 300 m wide, with sharp contacts, locally with chilled margins of porphyritic texture. The edge of the body is highly sheared, and the interior is cut by several shear zones. This is a muscovite-biotite monzogranite, with garnet, titanite, and zircon as accessories. A contact aureole may be indicated by relicts of porphyroblasts of biotite. Intruding the elongated tonalite of E1 Realito, Brogioni et al. (1994) also recognized a NEelongated stock of granite, associated with aplitic dykes, which was later deformed together with the tonalitic country rock. It is a mediumgrained garnet-bearing monzogranite, with incipient foliation defined by orientation of muscovite, biotite and quartz aggregates. Quartz, microcline, and plagioclase are incipiently recrystallized. Chemical characteristics. For the Rio Quinto granite only two samples were chemically analysed for major elements. These have 72-74% SiO2, low K20/Na20 ratio (0.8-0.9), and a slightly peraluminous character (A/CNK = 1.13-1.17), in accordance with minor muscovite

The size and shape of the La Escalerilla pluton is distinctly different from the other plutons of the Sierra de San Luis. It is a thin pluton, c. 52 km long and about 2-10 km wide, with an inflexion in the central part (Fig. 1). The pluton has been described to the north of this inflexion. Despite the later deformation which converted the country rocks to phyllites, the intrusive character and contact aureole can be recognised at the western border, where the pluton was intruded into still undeformed sedimentary rocks (von Gosen & Prozzi 1996; von Gosen in press). Some granitic dykes within these sedimentary rocks are related to the pluton. To the south, the pluton, as well as aplitic dykes related to it, cut the eastern border of Las Verbenas tonalitic pluton (Sato 1993) and were later deformed together with the country rocks. On the eastern side of the La Escalerilla pluton, the mica-schist belt is in contact with the granite along a mylonite zone affecting both rock units (Llano et al. 1987; von Gosen & Prozzi 1996, this volume). The northern part of the pluton has a ductile deformation with a NNE-SSW trending foliation (von Gosen & Prozzi 1996), which is more intense at the eastern border. The La Escalerilla pluton is of monzogranitic composition, varying to granodiorite. Textures are equigranular to inequigranular, with K-feldspar megacrysts. Inside the pluton aplitic syn-plutonic dykes with pegmatitic cores are common. They are related to the late stages of crystallization of the pluton, and emplacement was favoured by volume contraction due to solidification. These syn-plutonic dykes suggest

GRANITOIDS OF THE SIERRA DE SAN LUIS

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Fig. 5. (a-e) R1-R2 diagrams with discriminant tectonic fields after Batchelor & Bowden (1985) and (d-f) chondrite-normalized REE (Taylor & McLennan 1985) of the Famatinian granitoids from the Sierra de San Luis grouped according to their relative age respect to the main orogenic phase (Ordovician age). (a) and (d) Early Famatinian deformed granitoids prior to the main orogenic phase; (b) and (e) typical Famatinian granitoids related to the early stages of the main orogenic phase; (c) and (f) Late Famatinian, post-orogenic granitoids. that the country rock of the pluton was rigid enough to prevent its collapse. Plagioclase is weakly zoned with a composition ranging from An30 in the rare grano-

dioritic facies to An20 in the most abundant monzogranites. Myrmekitic textures are common. K-feldspar is commonly perthitic and the megacrysts are up to 8cm long. Biotite is

332

E. J. LLAMBIAS E T AL.

variable in amount, ranging up to 14%, but mainly 1-6%. Muscovite is rare (1-3%) and probably not primary in origin. Age. A Rb-Sr errorchron date of 414Ma with an initial 87Sr/86Sr of 0.7106 ( M S W D = 6.3) is delineated by five samples (Sato, unpublished data). Two of these samples were taken from the eastern mylonite zone and have the highest 87Rb/86Sr values (close to 40). They define the slope (and hence the age) of the errorchron, the 87Rb/86Sr values of the remaining three samples being less than 6. Therefore, this age may be interpreted as the time of mylonitization rather than that of the crystallization of the pluton. According to yon Gosen & Prozzi (1996), mylonitization is coeval with the main deformation of the pluton. However, a similar age of 403 + 6 Ma (Sims et al. this volume) was also obtained by the SHRIMP U-Pb zircon method. Chemical characteristics. The few samples analysed indicate a SiO2 range of 68-76%, a weak peraluminous character ( A / C N K = 1.02-1.22), and a high K20/Na20 ratio (1.2-1.7). In the R1R2 diagram (Fig. 5a), the samples plot in the syn-collisional field, close to those of the earlyto syn-orogenic garnet-muscovite granodiorites and granites. The only sample analysed for REE contents displays a negative Eu anomaly, suggesting fractionation of feldspar. The La/Yb ratio of one sample is 29 and is higher than those of the Early Famatinian tonalite group, suggesting a change in the source or an increase in the thickness of the crust. Discussion. The

La Escalerilla pluton is a unique deformed pluton, of large size and high aspect ratio. We consider that the latter is an original characteristic and not due to stretching during deformation. Had this been the case, the magnitude of the deformation would have been very great and this is not the case. If our assumption is correct, emplacement could have been controlled by a fracture. In addition, some characteristics of the intrusive body, such as the discordant contacts, granitic dykes in the country rock, and the internal syn-plutonic aplitic dykes, suggest a rigid behaviour of the country rock during emplacement. Some evidence, e.g., the deformation of the granite dykes in the country rock and the solid deformation of the granite itself (see yon Gosen & Prozzi 1996. this volume), indicate that intrusion preceded the metamorphism and deformation of the phyllite group (von Gosen & Prozzi 1996). The Rb-Sr and U-Pb ages mentioned above are younger

than the Rb-Sr errorchron reported for early- to syn-orogenic granitoids of the Paso del Rey-La Carpa area (Llambias et al. 1991). As discussed below, both emplacement and deformation of the latter granitoids are interpreted as having occurred within a short time span at about 454Ma. Hence, the ages of 414-403Ma obtained for the La Escalerilla granite seem to be too young either for crystallization or deformation of the pluton, assuming that deformation was contemporaneous in both granitoid groups. Moreover, these ages are also close to the Rb-Sr isochron age of 408 + 25 Ma for the oldest pluton of the undeformed Las Chacras batholith (Brogioni 1993). Thus, the timing of emplacement of the La Escalerilla granite remains unsolved with presently available data (for a more detailed discussion see Von Gosen & Prozzi this volume).

Garnet-biotite--muscovite granodiorites and granites: Ordovician early to syn-orogenic Famatinian granitoids These garnet-biotite-muscovite granodiorites and granites are located in the central part of Sierra de San Luis, together with the mica-schist and the adjacent biotite-garnet gneisses. The intrusive bodies are of small size, commonly less than 2 km 2. The best known areas are those of Paso del Rey, Cruz de Carla, Rio de La Carpa, and Cerros Largos. In the northeastern part of Sierra de San Luis, the La Ci6naga and La Represa plutons have also been recognized (Ortiz Su/trez 1996), as well as the small lensoid granitoid group of La Tapera near the eastern border (L6pez de Luchi 1987). All these granitoids are closely related in space and time with swarms of quartz-feldspar-muscovite pegmatite dykes, some of which contain beryllithium-tin deposits (Ortiz Su/trez & Sosa 1991). The composition ranges from biotite tonalite to muscovite-garnet granite (Fig. 4b). Pegmatitic segregations with diffuse borders are abundant. The contacts with the country rock are sharp and parallel to the metamorphic foliation. In many cases, such as in the Paso del Rey and Rio de la Carpa plutons, the edge of the igneous body is folded together with the country rock. The internal foliation of the igneous bodies is more intense towards the border zones, being parallel to the regional metamorphic foliation, which normally wraps around the bodies but also cross-cuts them along shear zones. Most of the plutons are elongated, with high aspect ratios (up to 18:1) and major axes

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medium grained. Grain boundaries are lobate to ameboidal. Plagioclase (An20) is weakly zoned with antiperthitic cores and albitic rims. Inclusions of muscovite, biotite, and quartz are common. Microcline contains inclusions of quartz and plagioclase with albitic rims and often replaces plagioclase. Myrmekite is abundant. Individual crystals or aggregates of biotite are orientated coincident with the regional foliation. Muscovite appears as individual crystals, often showing symplectitic intergrowth with quartz and plagioclase. The content of muscovite and the muscovite/biotite ratio increase with microcline content. Garnet is sub- to anhedral (1-2 ram). Fibrolitic sillimanite has been found in the granitic sheets north of Paso del Rey. For the Paso del Rey and Rio de la Carpa bodies Llambias et al. (1991) obtained a ten-point errorchron of 454~21 Ma, with an initial 87Sr/86Sr 0.7118+0.0001 (McIntyre model III), with MSWD = 12.8. These authors related this age to the time of emplacement, but in view of the recrystallization that occurred during the metamorphism we cannot rule out the possibility that the age is related to this event. Nevertheless, the chemical characteristics of these igneous rocks, typical of anatectic magmas, and the high initial 87Sr/86Sr ratio, lead us to believe that emplacement, deformation and metamorphism were contemporaneous within the error of the errorchron. K - A r dates on biotite (Varela et al. 1994) range 372-392 Ma, suggesting that either very slow cooling or a thermal event reset the K - A r system at this time. The latter possibility cannot be ruled out, since some post-orogenic granites (e.g. El Morro and La Mesilla) have similar ages. Another Rb/Sr whole rock age of 485 + 30 Ma (initial 87Sr/86Sr ratio = 0.7073) was reported by Ldpez de Luchi (1987) for the La Tapera pluton in the eastern part of the Sierra de San Luis. Age.

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parallel to the regional structures, but their shapes are variable; some are tightly folded (Fig. 6), others are elongated and boudinaged (Fig. 7), and still others are elliptical (Fig. 8). Sheet-like bodies, 1-5 m thick and parallel to the foliation, are also common. Disc-shaped porphyroblasts (andalusite?) about 1 cm in diameter, in the country rock, spatially related to the Paso del Rey and Rio de la Carpa intrusive bodies, are highly sheared and retrograded to a mixture of sericite and quartz (von Gosen 1998; von Gosen & Prozzi this volume). On the other hand, the intrusion of the plutons of the northeastern part of the Sierra de San Luis (La Represa and La Ci6naga granodiorites and granites), was associated with a lowpressure regional metamorphism, with growth of porphyroblasts ofandalusite (Ortiz Sufirez 1996). The garnet-biotite-muscovite granodiorites and granites are light grey, commonly fine to

Chemical characteristics. Although most of the analysed rocks are granodiorites according to the QAP diagram (Fig. 4b), the samples from Paso del Rey and Rio de La Carpa plot in the granite field of the chemical classification of Middlemost (1994). The SiO2 content varies from 67% to 75%. The intrusive rocks of the Cruz de Carla area have lower SiO2 contents (67-70%) than those of the Paso del Rey-Rio de La Carpa area (70-75%). In the R1-R2 diagram (Fig. 5b), most rocks plot in the syn-collision field, although those of Cruz de Carla plot closer to the field of pre-plate collision and mantle fractionates. However, no differences have been

334

E. J. LLAMBIAS E T AL. absence of a negative Eu anomaly indicate poorly differentiated magmas. The water was probably derived from protoliths rich in hydrated minerals such as micas (Holtz & Johannes 1994). This fact, added to the high initial 87Sr/86Sr ratio (0.7118, Llambias et al. 1991), supports the idea of magmas derived from crustal rocks. The high La/Yb ratios (15-37) suggest the presence of garnet in the source, which can be interpreted as the result of a thickening of the crust during the evolution of these igneous bodies. The small volume of the magma indicates a low melt proportion, characteristic of some anatectic magmas. The emplacement of the garnetbiotite-muscovite granodiorites has been described as syn-kinematic with respect to the Ordovician metamorphism and deformation (Llambias et al. 1991; Ortiz Sufirez et al. 1992; Llambias et al. 1996a), on the basis of the small size of the bodies, their low rheological contrast with the country rock, their high aspect ratio, and the parallelism of the main axis with the foliation of the country rock. However, the highly sheared edges of most of the igneous bodies, with decreasing intensity towards the interior of the plutons, and their shapes, suggest that they behaved as rigid bodies during deformation (see von Gosen & Prozzi this volume). The high initial 87Sr/g6Sr ratio, the chemical signature typical of low temperature anatectic magmas, and the uplift of the crust that followed the main orogenic phase, all suggest that the crust was either already thickened or was actively thickening during the evolution and emplacement of these granites. For these reasons, we tentatively correlate these granitoids with the first stages of deformation during the main orogenic phase of the Famatinian cycle. Discussion.

Fig. 7. Geological map of Rio de La Carpa pluton.

observed in the K20/Na20 ratio (0.6-1.5; 10 of 17 analysis are 1.1). These values represent a metaluminous to slightly peraluminous signature. The Rb/Sr ratios (0.501.08) are low and may be correlated with the abundance of plagioclase, and the lack of a negative Eu-anomaly (Fig. 5e). The abundance of late magmatic pegmatitic segregations within these granitic rocks, and the close time and spatial relationships to the pegmatite swarms in the country rocks, suggest that the water content of the magmas was probably close to saturation. This conclusion is also supported by compositions that approach the saturated haplogranite minimum melt (Tuttle & Bowen 1958; Luth et al. 1964), especially in the case of the Paso del Rey-Rio de La Carpa area. Nevertheless, this high water content cannot be explained by enrichment due to magmatic differentiation, because the low Rb/Sr ratios and the

Biotite, biotite-hornblende granites and tonalites: Devonian to Early Carboniferous, Late Famatinian, post-orogenic granitoids Biotite granites and biotite-hornblende granites are widely distributed in the Sierra de San Luis, as isolated plutons or as batholiths. Less widespread tonalitic plutons are spatially related to the granites and have similar relationships with the metamorphic country rock. The largest outcrops correspond to the Las Chacras batholith (Fig. 1), of around 800km 2, which is composed of four major monzogranitic plutons of circular section aligned in a NW-SE trend (Brogioni 1991). Mutually intrusive relations and the lack of textural variations along the

GRANITOIDS OF THE SIERRA DE SAN LUIS

Fig. 8. Geological map of the Paso del Rey pluton.

border zones make it difficult to recognise the contacts. Other circular plutons are the isolated Renca stock (L6pez de Luchi 1993), emplaced at the southeastern continuation of the Las Chacras trend, and the La Totora pluton, whereas the E1 Morro pluton has an oval shape. To the north and northeast of the Las Chacras batholith there are several monzogranitic plutons, e.g., the E1 Telarillo, E1 Hornito, Los Alanices, and E1 Pefi6n Grande plutons (Fig. 1), and three with tonalitic composition. The Rodeo Viejo (30 km2), E1 Salado (3.5 km 2) and La Cienaguita plutons have a predominantly hornblende-biotite tonalitic composition (Ortiz Sufirez 1996). The first of these has an elongated shape with its major axis NW-SE, oblique to the metamorphic foliation. The contacts are sharp, with a narrow chilled border. The contact aureole consists of hornfels with garnet, sillimanite, andalusite and staurolite. Xenoliths and small pendants of the metamorphic country rock indicate postmetamorphic emplacement. The igneous bodies mostly cross-cut the different metamorphic rocks and their foliation; only locally are they concordant with the

335

foliation of the mica-schists, e.g., in a zone of up to 600 m wide adjacent to the largest pluton (c. 300 km 2) of the Las Chacras batholith. In the E1 Morro pluton several pendants of metamorphic country rock have been described (Llambias & Malvicini 1982; Quenardelle 1993). They have angular and planar borders and are hundreds of metres long. The variable orientation of the metamorphic foliations within these pendants indicates that they were rotated during intrusion. Chilled margins are often observed, composed of fine-grained monzogranites, aplites and porphyritic rhyolites. Aplitic dykes are commonly related to the igneous bodies. They cut the country rock discordantly; in the Las Chacras batholith some of the aplitic dykes have an annular pattern. The monzogranites have equigranular to inequigranular textures, with perthitic K-feldspar megacrysts up to 12 cm long, containing inclusions of euhedral and zoned plagioclase, quartz and biotite. Plagioclase (An20-30) grains outside the K-feldspar megacrysts are commonly anhedral with sutured borders. Biotite and amphibole are common, whilst magnetite and subordinate ilmenite are always present. Titanite (<2mm) and apatite (<6 mm) are typically large; allanite is rare. Mafic microgranular enclaves of variable composition, mainly tonalite, mafic monzonite and also syenite, are abundant (Brogioni 1991; Quenardelle 1993; L6pez de Luchi 1993). The tonalitic rocks of Rodeo Viejo consists of hornblende, biotite, plagioclase (An34-48) and quartz (Ortiz Sufirez 1996). Age. This group of granitoids is clearly younger with respect to deformation and metamorphism. The K-Ar and Rb-Sr ages range between 408 and 320 Ma, suggests a long period of post-orogenic intrusion. The reported ages are: 408 + 25 Ma, Rb-Sr isochron for the La Mesilla pluton of the Las Chacras batholith (Brogioni 1993); 382 + 4 M a , Rb-Sr errorchron for the E1 Morro pluton (Varela et al. 1994); 380-360 Ma, K-Ar data on biotite and amphibole from the E1 Morro pluton (Lema 1980); 336-320Ma, K-Ar data for the Las Chacras and Potrerillo plutons from the Las Chacras batholith (Brogioni 1987). The two Rb-Sr isochrons reported by Brogioni (1993) from different rock facies of the Las Chacras pluton have large errors (317+54 Ma, 307• although MSWD is low in both cases. K-Ar ages of 301-305Ma for the La Totora pluton (Kittl 1965) are considered to be too young: the basement rocks would have been already exhumed at that time, since they have an Upper Carboniferous sedimentary cover

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(e.g., Bajo de Veliz Formation in the Sierra de San Luis, Azcuy et al. 1987). A SHRIMP U-Pb zircon age for the Renca pluton gave 393 -t- 5 Ma (Sims et al. this volume). Chemical characteristics. Silicic rocks predominate with an SiO2 range of 61-76%, mostly between 66% and 76%. The modal composition, characterized by hornblende and titanite, indicates a metaluminous character. The chemical composition of the Las Chacras batholith and the E1 Morro pluton (Brogioni 1991; Quenardelle 1995) show a metaluminous signature transitional to peraluminous (A/CNK = 0.84-1.20). In the E1 Morro pluton the K20/Na/O ratio ranges from 1.44 to 1.96, explaining the early crystallization of K-feldspar megacrysts (Holtz & Johannes 1991). This is also supported by the magmatic orientation of the K-feldspar megacrysts, indicating that they grew before reaching the critical volume fraction of crystals, above which the effective viscosity is so high that the magma would behave as a solid. Rb/Sr ratios are low (0.5-1.0) indicating a lack of feldspar fractionation; this is corroborated by the very weak negative Eu anomalies of two samples analysed from the El Morro pluton (Quenardelle 1995). The La/Yb ratios (20 and 23) of these samples are lower than those of the garnetmuscovite granodiorites (Fig. 5f), indicating an important change in the source. The origin of the granitic magmas has been explained as a consequence of thermal input from a mafic underplate (Brogioni 1991; Quenardelle 1995), and in Renca by decompression (Lopez de Luchi 1996). The low S7Sr/86Sr ratios (0.7036-0.7053) for the Las Chacras batholith support the involvement of magma generated in the mantle (Brogioni 1993). On the other hand, the higher initial 87Sr/86Sr ratio of 0.706 obtained at E1 Morro pluton (Varela et al. 1994), indicates some crustal contamination. Discussion. The relationship of the intrusive bodies with the country rock indicates generally post-orogenic emplacement. The predominance of granites over granodiorites and tonalites is more typical of granitic or silicic suites than of the calc-alkaline suites of magmatic arcs. The less common tonalitic and granodioritic bodies, and the abundance of mafic microgranular enclaves, is consistent with the mafic underplating model. The large volume of monzogranitic rocks is clear evidence of a high degree of melting in the source, a condition which favoured the emplacement of the large plutons in the upper crust.

Magmatic evolution Emplacement of the Early Famatinian tonalitic intrusives occurred at upper crustal levels (Llambias et al. 1996b) within a country rock composed of already deformed gneisses and unmetamorphosed sedimentary rocks, the latter being transformed into phyllites by subsequent deformation (von Gosen & Prozzi 1996, this volume). The previous metamorphic history and age of the gneissic rocks are unknown. Uplift of these deeper levels may be attributed to the end of a pre-Famatinian orogenic cycle (Llambias et al. 1996b; Sato et al. 1996). According to Prozzi (1990), the sedimentary rocks have lithological characteristics similar to the Upper Proterozoic-Lower Cambrian Puncoviscana Formation of NW Argentina. The geological and chemical features of the tonalites agree with the evolution of a magmatic arc related to subduction of oceanic lithosphere beneath a continent. The rather high initial S7Sr/S6Sr ratio of these rocks (c. 0.708) suggests some kind of crustal contamination, a process possible during ascent of the magma through reworked crust. The parallelism between the NNE-SSW distribution of the tonalites (Fig. 1) and the foliation trend of the subsequently deformed rocks may suggest continuous evolution of a single orogenic process, supporting the Early Famatinian age of this arc (Sato et al. 1996) Rocks of similar age and composition crop out 200-300 km to the NNW of the Sierra de San Luis, along the eastern margin of Sierra de la Huerta-Valle F6rtil (Mirr6 1971; Toubes Spinelli 1983; Castro de Machuca et al. 1996; Vujovich et al. 1996) and the Sierra de Chepes and Las Minas (Dahlquist & Baldo 1996; Pankhurst et al. and Sims et al. this volume), which according to Caminos (1972) are of batholithic scale. Although some Rb-Sr ages for the Chepes granodiorite are slightly younger (Pankhurst et al. 1996), we do not exclude the possibility that these rocks correspond to the northern continuation of the magmatic arc defined for Sierra de San Luis. The garnet-muscovite granodiorites and granites differ from the tonalitic suite in composition, and shape and size of the igneous bodies, indicating an important change in tectonic setting. The higher initial 87Sr/86Sr ratios and S-type compositions suggest a crustal magma source close to the level of emplacement. The small size of the igneous bodies indicate that the magma failed to coalesce into larger bodies. A possible reason for this failure may be related to derivation from a protolith rich in muscovite; according to Holtz & Johannes (1994), magmas formed by dehydration melting at low

GRANITOIDS OF THE SIERRA DE SAN LUIS temperature involving muscovite breakdown can only produce low melt fractions. Although extraction and coalescence is favoured by the low viscosity, the small volume of the melts and the HzO-rich conditions make it difficult for the magma to ascend. The above mentioned features, typical of anatectic magmas, favour the model of magma genesis in a thickened crust, which is also supported by the high La/Yb ratios. In this model, the crust would have been thickened during the period between development of the magmatic arc and the emplacement of the granites. Taking into account that several of these bodies have structural features indicating rigidity during the main orogenic phase, it is possible to infer that magma generation

occurred either at the beginning of, or simultaneously with, crustal thickening. The compression in the Sierra de San Luis may be the result of the collision of an allochthonous terrane (recently interpreted as a Laurentia-derived terrane) during the Early Palaeozoic, as proposed by several authors (Ramos 1988; Dalla Salda et al. 1992; Benedetto 1993; Astini et al. 1995; Keller & Dickerson 1996). After the main orogenic phase, there followed a period of isostatic adjustment, with general regional uplift and probable decrease in crustal thickness. During this period the large granitic batholiths were emplaced. Their intrusion along fractures, such as that which controlled the distribution of the plutons in the Las Chacras batholith (Fig. 1), and the common circular

compression (collision?)

subduction magmatic arc

337

?

extension

early-orogenic Post-orogenic anatectic transitional to granites anorogenic thickening of the crust

Famatinian

cycle

f

187Sr/SSSr0 La/Yb

0.711

30 0.709 20 0.707 10 0.705

Tonalite group

La Escalerilla granite

UpperOrdovician Cambrian Lower

9

gt-m..sgranites

bio-hblgranites Upper Carboniferous Silurian to O r d o v i c i a nLower

Fig. 9. Cartoon summarising the evolution of the Famatinian granitoids of the Sierra de San Luis. The blocks with vertical lines show the range of the La/Yb ratios. For discussion see text.

338

E. J. LLAMBIAS E T A L .

sections of the plutons, with internal concentric zoning and annular dykes (Brogioni 1991), are typical of post-orogenic to anorogenic suites emplaced in rigid and cold upper crust. Some of the plutons probably represent the roots of volcanic centres. Most of the rocks have some attributes of hypersolvus granites (high-temperature and volatile-poor), a characteristic which favoured ascent of the magma to the upper crust. Between the garnet-biotite granodiorites and granites dated at 454 Ma and the oldest plutons of the post-orogenic suite, dated at 408 Ma, there is a period of about 45 Ma. During this period, an important change occurred in the condition of magma generation, resulting in increased melt production. The possible scenario is a collision, as mentioned above, with transient cooling of the crust followed by decompression, as has been proposed for the Renca ptuton (Lopez de Luchi 1996), and thermal input caused by the formation of a mafic underplate at the base of the crust (Brogioni 1993; Quenardelle 1995). The age and composition of the post-orogenic intrusive rocks of the Sierra de San Luis are similar to those of others in the Eastern Sierras Pampeanas, e.g., the Achala (Lira & Kirschbaum 1990; Rapela et al. 1991) and Cerro Aspero (Pinotti et al. 1996) batholiths in the Sierras de C6rdoba. This means that, despite the differences in the previous metamorphic history between San Luis and C6rdoba, the evolution of Late Famatinian granitoids was a common feature to both. In Fig. 9 we summarize schematically the igneous evolution of the Sierra de San Luis, relating the different La/Yb ratios to initial 87Srff6Sr ratios. The early Famatinian tonalites have low 87Sr/86Sr and La/Yb ratios, characteristic of magmatic arcs. Both ratios increased in the La Escalerilla granite and in the Ordovician granitoids, probably due to thickening of the crust. After this, the crust started to thin and the melt production significantly increased; initial 87Sr/86Sr ratios became lower, suggesting a source in the mantle.

Conclusions The Famatinian cycle is well represented in the Sierra de San Luis. Remnants of an older orogenic cycle have been recognized, but they are poorly constrained. The Late Cambrian to Early Ordovician deformed tonalitic rocks are part of an Early Famatinian magmatic arc, emplaced into shallow levels of the crust. The geological relations and age of the La Escalerilla granite are as yet unclear.

The deformed garnet-biotite-muscovite granodiorites, granites and related pegmatites of Ordovician age represent largely anatectic magmas related to a thickened crust. During this time melt production was relatively low and emplacement was close to the magmatic source. The setting is early to syn-orogenic and probably related to a collision. The undeformed biotite and biotite-hornblende granites of Devonian to Early Carboniferous age are post-orogenic, related to uplift. Melt production was very high, probably caused by mafic underplating. Financial support from CONICET and Universidad Nacional de San Luis allowed this study to be carried out. We thank W. von Gosen (Erlangen University) for many discussions during field trips about the relationship of the granitoids with the metamorphic and deformed rocks, and also for critical review of the manuscript. We thank R. Pankhurst for many comments on the drafts. This manuscript benefited from critical reviews by D. Hutton and an anonymous reviewer. The English version was greatly improved by Anne Dalton.

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, SATO, A. M., PROZZI, C. & SANCHEZ, V. 1996b. Los pendants de gneises en el plut6n Gasparillo: evidencia de un metamorfismo pre-Famatiniano en las sierras de San Luis. XIII Congreso Geoldgico Argentino y III Congreso de ExploraciOn de Hidrocarburos, Buenos Aires, Actas, V, 369-376. LLANO, J., CASTRO DE MACHUCA, B., ROSSA, N. & VACA, A. 1987. Las rocas catacl/tsticas en el perfil Valle de Pancanta, Paso del Rey, Sierra de San Luis, Repfiblica Argentina. X Congreso Geoldgico Argentino, Tucumfin, Actas, III, 31-34. LOPEZDE LUCHI, M. 1987. Caracterizacidn geolbgica y geoquimica del plutdn La Tapera y el batolito de Renca, provincia de San Luis. X Congreso Geoldgico Argentino, Tucum/m, Actas, IV, 84-87. - - 1 9 9 3 . Caracterizaci6n geoldgica y emplazamiento del Batolito de Renca. XII Congreso Geoldgico Argentino y H Congreso de Exploracidn de Hidrocarburos, Mendoza, Actas, IV, 42-53. - - 1 9 9 6 . Enclaves en un batolito postect6nico: petrologia de los enclaves microgranulares del batolito de Renca, Sierras Pampeanas, San Luis. Revista de la Asociaeidn Geoldgica Argentina, 51, 131-146. LUTH, C., JAHNS, R. H. & TUTTLE, O. F. 1964. The granitic system at pressure of 4 to l0 kilobars. Journal of Geophysical Research, 69, 759-773. MIDDLEMOST, A. K. 1994. Naming materials in the magma/igneous rock system. Earth Science Reviews, 37, 215-224. MIRRI~, J. C. 1971. Caracterizacibn de una comarca de metamorfismo regional de alto grado: la Sierra de Valle F6rtil, provincia de San Juan. Revista de la Asociacidn Geoldgica Argentina, 26, 113-127. ORTIZ SUAREZ, A. 1996. Geologia y petrografia de los intrusivos de Las Aguadas, provincia de San Luis. Revista de la Asociaeidn Geoldgica Argentina, 51, 321-330. -& SOSA, G. 1991. Relaciones gen+ticas entre las pegmatitas portadoras de estafio y metamorfitas asociadas en la zona de La Carolina/San Francisco del Monte de Oro. Provincia de San Luis. Revista de la Asociacidn Geoldgica Argentina, 46, 339-343. , PROZZI, C. & LLAMBiAS,E. J. 1992. Geologia de la parte sur de la Sierra de San Luis y granitoides asociados, Argentina. Estudios Geol6gicos, 48, 269-277. PANKHURST, R. J., RAPELA, C. W., SAAVEDRA, J., BALDO, E., DAHLQUIST, J. & PASCUA, I. 1996. Sierra de Los Llanos, Malanzan and Chepes: Ordovician I- and S-type granitic magmatism in the Famatinian orogen. XIII Congreso Geoldgico Argentino y III Congreso de Exploracidn de Hidrocarburos, Buenos Aires, Actas, V, 415. & FANNING, C. M. ]998. "]'he Famatinian magmatic arc in the central Sierras Pampeanas. This volume. PINOTTI, L. P., CONIGLIO, J. E. & LLAMBiAS, E. J. 1996. Caracteristicas geol6gico-estructurales del plut6n Alpa Corral, 32~176 y 64~ 64~ Sierras Pampeanas de C6rdoba, Argentina. XIII Congreso Geol6gico Argentino y 1II Congreso de Exploracidn de Hidrocarburos, Buenos Aires, Actas, III, 477-486.

PROZZI, C. 1990. Consideraciones acerca del basamento de San Luis, Argentina. XI Congreso Geol6gico Argentino, San Juan, Actas, I, 452-455. QUENARDELLE, S. M. 1993. Caracterizaci6n geol6gicapetrol6gica del Granito San Jos6 del Morro, provincia de San Luis. XII Congreso Geol6gico Argentino y Segundo Congreso de Exploraci6n de Hidrocarburos, Mendoza, Actas, IV, 61-67. 1995. Petrografia y geoquimica del plut6n San Jos6 del Morro, provincia de San Luis. Revista de la Asociaci6n Geol6gica Argentina, 50, 229-236. RAMOS, V. A. 1988. Late Proterozoic-early Paleozoic of South America: A collisional history. Episodes, 11, 168-175. RAPELA, C. W., PANKHURST, R. J., CASQUET, C., BALDO, E., SAAVEDRA,J., GALINDO, C. & FANMNG, C. M. 1998. The Pampean Orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de C6rdoba. This volume. --, KIRSCHBAUM,A. & BALDO, E. G. A. 1991. Facies intrusivas de edad carbonifera en el batolito de Achala: evidencia de una anatexis regional en Sierras Pampeanas? 6~Congreso Geoldgico Chileno, Vifia del Mar, Resumenes Expandidos, 40-43. SANCHEZ, g., ORTIZ SUAREZ, A. & PROZZI, C. 1996. Geologia y petrografia de la tonalita Bemberg, provincia de San Luis. XIII Congreso Geoldgico Argentino y III Congreso de Exploracidn de Hidrocarburos, Buenos Aires, Actas, III, 669-677. SATO, A. M. 1993. Deformaci6n de la tonalita precinem/ttica de Las Verbenas, Sierra de San Luis, Argentina. Primer Simposio Internacional del Neoproterozoico-Cdmbrico de la Cuenca del Plata. DINAMIGE-Facultad de Agronomia, Montevideo, 1, NE25. , ORTIZ SU.~.REZ,A., LLAMBIAS,E. J., CAVAROZZI, C. E., S~i.NCHEZ,g., VARELA, R. & PROZZI, C. 1996. Los plutones pre-Ocl6yicos del sur de la Sierra de San Luis: arco magmfitico al inicio del ciclo Famatiniano. XIII Congreso Geoldgico Argentino y III Congreso de Exploracidn de Hidrocarburos, Buenos Aires, Actas, V, 259-272. SIMS, J. P., IRELAND,T. R., CAMACHO,A., LYONS, P., PIETERS, P. E., SKIRROW, R. G., STUART-SMITH, P. G. & Mm6, R. 1998. U-Pb, Th-Pb and Ar-Ar geochronology from the southern Sierras Pampeanas, Argentina: implications for the Palaeozoic tectonic evolution of the western Gondwana margin. This volume. TAYLOR, S. R. & MCLENNAN, S. M. 1985. The continental crust." its composition and evolution. Blackwell, Oxford. TOUBES SPINELLI,R. O. 1983. Edades potasio argdn de algunas rocas de la sierra de Valle F6rtil, provincia de San Juan. Revista de la Asociaci6n Geol6gica Argentina, 38, 405-411. TROMPETTE, R. 1994. Geology of western Gondwana (2000-500 Ma). A. A. Balkema, Rotterdam. TURNER, J. C. M. & MI~NDEZ, V. 1975. Geologia del sector oriental de los departamentos de Santa Victoria e Iruya, provincia de Salta, Repfiblica Argentina. Academia Nacional de Ciencias, Boletin, 51, 11-24.

G R A N I T O I D S OF T H E S I E R R A DE SAN LUIS & MI~NDEZ,V. 1979. Puna. In: TURNER, J. C. M. (ed.) Segundo Simposio Geologia Regional Argentina. Academia Nacional de Ciencias, Cbrdoba, 13-56. TUTTLE, O. F. & BOWEN,N. L. 1958. Origin of granite in the light of experimental studies in the system NaAISi3Os-KAISi3Os-SiOe-H20. Geological Society of America, Memoirs, 74. VARELA, R., LLAMBiAS,E. J., CINGOLANI, C. & SATO, A. M. 1994. Datacibn de algunos granitoides de la Sierra de San Luis (Argentina) e interpretaci6n evolutiva. 7~ Congreso Geoldgico Chileno, Concepcion, Actas, 2, 1249-1253. --, LOPEZ DE LUCHI, M., CINGOLANI, C. & DALLA SALDA, L. 1996. Geocronologia de gneises y granitoides de la sierra de Umango, La Rioja. Implicancias tect6nicas. XIII Congreso Geoldgico Argentino y III Congreso de Exploracidn de Hidrocarburos, Buenos Aires, Actas, IIl, 519-527. VON GOSEN, W. 1998. The phyllite and micaschist group with associated intrusions in the Sierra de San Luis (Sierras Pampeanas/Argentina) - structural and metamorphic relations. Journal of South American Earth Sciences, 11, 79-109.

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Intrusions, metamorphism and deformation with transpressive effects in the southwestern part of the Sierra de San Luis (Sierras Pampeanas, Argentina). Journal of South American Earth Sciences, in press. & PROZZl, C. 1996. Geology, structure, and metamorphism in the area south of La Carolina (Sierra de San Luis, Argentina). XIII Congreso Geol6gico Argentino y III Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, I1, 301-314. & 1998. Structural evolution of the Sierras de San Luis (eastern Sierras Pampeanas, Argentina): implications for the proto-Andean margin of Gondwana. This volume. VUJOVICH, G., GODEAS, M., MARLY, G. & PEZZUTTI, N. 1996. E1 complejo magmfitico de sierra de La Huerta, provincia de San Juan. XIII Congreso Geoldgico Argentino y III Congreso de Exploracidn de Hidrocarburos, Buenos Aires, Actas, III, 465-475. ZARDINI, R. 1966. Composici6n, estructura y origen del plutbn de la Pampa del Tamboreo (provincia de San Luis). Acta Geoldgica Lilloana, 8, 50-73.

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The Famatinian magmatic arc in the central Sierras Pampeanas: an Early to Mid-Ordovician continental arc on the Gondwana margin R. J. P A N K H U R S T

1, C. W. R A P E L A 2, J. S A A V E D R A 3,

E. B A L D O 4, J. D A H L Q U I S T 2, I. P A S C U A 5 & C. M. F A N N I N G 6

1British Antarctic Survey, NIGL Isotope Geosciences Laboratory, Keyworth NG12 5GG, UK 2 Centro de Investigaciones Geol6gicas, Universidad Nacional de La Plata, Argentina 3 CSIC, Instituto de Agrobiologia y Recursos Naturales, Salamanca, Spain 4 Departamento de Geologia, Universidad de C6rdoba, Argentina 5 Departamento de Geologfa, Universidad de Salamanca, Spain 6 Research School of Earth Sciences, The Australian National University, Australia Abstract: A new multi-disciplinary study of the central Sierras Pampeanas encompasses

fieldwork, petrography, metamorphic and micro-structural analysis, geochemistry and geochronology. Remnants of a low-to-medium grade metasedimentary sequence, which also occurs in the Sierras de Cdrdoba to the east, are considered regionally equivalent to the Puncoviscana Formation; a ?mid-Cambrian Rb-Sr whole-rock isochron of 513 ~=31 Ma probably dates their main metamorphism. The predominant granitoids of the Los LlanosUlapes batholith constitute a calc-alkaline suite representative of the Famatinian subduction-related magmatic arc. The main granodiorite phase of the batholith is associated with an $2 fabric and shear zone formation, and was emplaced late during the deformational history of the metasediments. Conventional and SHRIMP U-Pb zircon dating yielded a combined age of 490 + 5 Ma. Younger monzogranites gave Rb-Sr wholerock ages of 470-450 Ma, typical of granites in the Sierra de Famatina, but geochemical continuity with the main granodiorite suite raises the possibility that these are partially reset ages. A minor cordierite granite phase is ascribed to local anatexis caused by heat from the granodiorites. All the calc-alkaline rocks of the Los Llanos-Ulapes batholith have high initial 87Sr/86Sr (0.7075-0.7105) and low cNdt (-4.6 to -6.3), inherited from lower crust. Sm-Nd model ages of 1600-1700Ma indicate that the underlying crust is identical to that beneath the foreland to the east. This part of the Famatinian arc was thus a continental magmatic arc and was established significantly before the arrival of the allochthonous Precordillera terrane in mid-Ordovician times. The rocks of the Sierras Pampeanas (Fig. 1) include widespread and varied granitoids of Lower Palaeozoic age. In the northern Sierras Pampeanas, notably in the Sierra de Famatina, the plutonic rocks and their volcanic counterparts have been identified as the Famatinian magmatic arc (see compilation by Acefiolaza et al. 1996). Although wide usage has probably resulted in the inclusion of rocks of different age, origin, and style of emplacement in this term, Famatinian igneous activity has recently been related to the approach and collision of an allochthonous terrane with continental South America in mid-Ordovician times (e.g., Datla Salda et al. 1992). This paper presents the results of a combined geological, geochronological and geochemical investigation in the central Sierras Pampeanas, as part of a complete E - W traverse carried out by Project ' P A M P R E ' (PampeanasPrecordillera). The results provide new evidence for the southward extension of the Famatinian arc and the circumstances of its evolution.

The sierras in the southern part of the province of La Rioja, at 30-32~ comprise a group of subdued mountainous ridges. The main ranges are known individually as the sierras of Los Llanos, Malanzfin, Chepes and, detached a little to the south, Ulapes (Fig. 1); unfortunately, there is no accepted collective name. They are located between a Cambrian basement block to the east in the Sierras de C6rdoba (Rapela et al. this volume) and the Sierra de Valle Fertil to the west, which has certain similarities to the study area but which has been considered to be partly allochthonous. The sierras of southern La Rioja are formed almost entirely of granitoid, with scarce outcrops of metamorphic rock representing remnants of the country rock into which the magmas were emplaced (Caminos 1979; Pascua et al. 1996; Dahlquist & Baldo 1996). Throughout the area, the granitoids exhibit strong mineralogical and geochemical affinities, and a uniform level of emplacement (Dahlquist &

PANKHURST, R. J., RAPELA, C. W., SAAVEDRA, J., BALDO, E., DAHLQUIST, J., PASCUA, I. & FANNING, C. M. 1998. The Famatinian magmatic arc in the central Sierras Pampeanas: an Early to Mid-Ordovician continental arc on the Gondwana margin. In: PANKHURST, R. J. & RAPELA, C. W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 343-367.

344

R. J. PANKHURST E T A L .

Fig. 1. Simplified sketch map of the Sierras Pampeanas. The metamorphic rocks and granitoids cover a wide range of types and ages (Proterozoic to Devonian). The main belt of the Famatinian magmatic arc (Early to Mid-Ordovician) is indicated by shading, and the principal sierras referred to in the text are identified by letters: F, Famatina; M, Malanzfin; L, Los Llanos; C, Chepes; U, Ulapes. Baldo 1996; Pascua et al. 1996). They have therefore been assigned batholithic status by Dahlquist & Rapela (1997) and are referred to here as the Los Llanos-Ulapes batholith. Our work was chiefly concentrated in the central part of this area, between Malanzfin and Olta, where detailed geological mapping was carried out. Although the study area has been previously considered as part of regional studies, this is the first detailed petrological and geochemical investigation. The area of the Andean hinterland between latitudes 26~ and 35~ (Fig. 1) has been the focus of recent terrane hypotheses concerning the early Palaeozoic geotectonic evolution of the proto-Andean continental margin (Ramos et al. 1986; Dalla Salda et al. 1992, Rapela et al. 1992; Astini et al. 1995; Ramos 1996). The timing and petrogenesis of the Los Llanos-Ulapes batholith has a significant bearing on these hypotheses,

especially insofar as the Famatinian arc is generally regarded as evidence of the subduction of ocean floor that accompanied the arrival of the allochthonous Precordillera terrane in midOrdovician times (e.g. Astini et al. 1995). The age of the main magmatic activity was first estimated as between 468 and 440 Ma by K - A r dating (Toselli & Gonzfilez 1978). However, the age, origin, petrogenesis and geotectonic setting of these rocks has not yet been clearly defined; this paper has, as equal objectives, the production of new data and the synthesis of research up to the present time. A new geological map (Fig. 2), new Rb-Sr and U - P b data and new petrological and geochemical data are presented for the Sierras of Chepes and Los Llanos. Preiminary Rb-Sr results, suggesting a Llandeilo-Caradoc age for some granites, were summarized by Pankhurst et al. (1996). Our conclusions are considered in relation to more

CENTRAL SIERRAS PAMPEANAS FAMATINIAN ARC

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Methods of investigation Detailed local geological mapping has been carried out by members of the project team

from the universities of C6rdoba and La Plata over the past three years. The full international team undertook field trips in 1994 and 1995, during which the majority of samples were collected. Analytical methods were essentially the same as in Rapela et al. (this volume), to which paper the reader is referred for a fuller description. Geochemical analyses for major and minor elements were obtained from ACTLABS, Canada (representative analyses are listed in Table 1). Geochronological analysis involved

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R. J. PANKHURST E T A L . A study of metamorphism and deformation in relation to granite emplacement (Dahlquist & Baldo 1996) was carried out along a profile in the NW sector of the Sierra de Chepes (Fig. 3). From the western end of this section, outcrops of the Chepes Granodiorite (see below) are followed by the metasedimentary sequence, increasing in grade from phyllite and spotted phyllite in the west, through porphyroblastic schist in the central part, to gneiss, migmatite and cordieritebearing granite in the east (corresponding to the elevated core of the sierra). In the west, the granodiorite shows deformation of variable intensity, with two main types of foliation: (1) predominantly N N E SSW, defined by mafic minerals and ellipsoidal microgranular enclaves, and (2) mylonitic foliation (N49~176 close to the contact with the metamorphic rocks. The mylonite belts (1-3 m) alternate with septa of phyllite over a distance of 100m. The first foliation is considered to be a result of primary magmatic flow, using the criteria of Paterson et al. (1989), but the orientation is of regional significance throughout the sierras of the study area, with a strike varying between NNE-SSW and NNW-SSE and a sub-vertical dip, predominantly to the E. The enclaves are invariably aligned in the plane of the foliation, with significant flattening perpendicular to it

the application of Rb-Sr, Sm-Nd and U-Pb isotopic techniques at the NERC Isotope Geosciences Laboratory, UK, and additional U-Pb zircon analyses by SHRIMP ion probe at ANU, Canberra. The results are reported in Tables 2-5.

Metamorphic complex Small, largely discontinuous, outcrops of metamorphic rock occur, principally in the margins and the central zone of the Los Llanos-Ulapes batholith (Fig. 2). They represent fragments and roof pendants of the country rock, and provide the only evidence for the nature of the basement into which the granitoids were emplaced. The metasedimentary sequence predominantly consists of metapelite, with alternating beds of siltstone and sandstone typically less than 1 m in thickness. The mineralogical composition is mainly quartz, plagioclase, K-feldspar, biotite and muscovite with cordierite, andalusite, opaque minerals and rare garnet. The regional metamorphic grade is mostly greenschist/amphibolite, and the protolith of this sequence would have been similar to that of the low-grade metapelites exposed along the western edge of the Sierra Grande de Cdrdoba (Rapela et al. this volume).

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CENTRAL SIERRAS PAMPEANAS F AM AT I NI AN ARC (aspect ratios of 3: 1). This suggests overall E - W contraction contemporaneous with emplacement of the igneous rocks. The mylonitic rocks close to the contact are of S - C type, with quartz aligned in veins or stringers that envelop the feldspars, signifying extreme deformation by plastic flow; the feldspar crystals show displacement on cleavage planes. No mafic minerals are seen but there is a concentration of Fe-Ti oxides in the C planes, together with fine granoblastic muscovite. Brittle-ductile conditions are indicated, with temperatures below the biotite stability limit in the greenschist facies. Farther west, in the western margin of the sierra, there is another mylonite-phyllonite belt, with mean strike and dip of N20~176 which shows evidence of more plastic conditions (coarser grain-size, biotite and hornblende markedly aligned with the quartz bands, and subhedral, rotated, feldspars). Micro-structural analysis shows that the metasedimentary rocks are polymetamorphic, with a relict low-grade paragenesis from a first, M l, metamorphic event, followed by growth of the porphyroblastic minerals (biotite, cordierite and andalusite) during a stage of low-pressure metamorphism, M2. Regional deformation produced a lepidoblastic matrix, resulting in a macroscopic schistosity, $2. The M2 event merits more detailed description on account of its regional significance to the emplacement of the granitoids. In the section A-A', the distribution of index minerals produced during M2 reveals a strong metamorphic zonation.

The phyllites have a penetrative foliation dipping mainly to the east, interpreted as $2 schistosity, syn-metamorphic and subparallel to SO (primary stratification). $2 dips at high, but variable, angles and strikes constantly N N E - S S W along the length of the profile. The typical mineral association is Bt + Ms + Qtz + Grt 4- Crd :t: Chl, using the nomenclature system of Kretz (1983), with tourmaline, zircon and opaque minerals as accessories. The biotite zone is restricted to the western border of the studied profile, but is well developed in other areas, e.g. to the west of Olta (Fig. 2). The metapelites contain porphyroblasts of Btl, surrounded by a lepidoblastic matrix of Bt2 and muscovite. The first prograde growth of small euhedral garnets (0.01 mm) is seen in these rocks. The otherwise restricted development of garnet suggests bulk composition control (i.e. only in MnO- and CaO-rich levels). The paragenesis of this zone

B i o t i t e zone.

349

is Bt + Ms + Qtz, and the most probable reactions are: Chl + Phengite + Qtz = Mt + Ms + H20, and Chl + Kfs = Bt + Ms. Cordierite zone. This is the widest zone, partly developed in the western sector of the phyllites, but principally in the schists of the central sector. It is characterized by the presence of porphyroblasts of poikilitic cordierite, of various dimensions (5-30mm) but generally becoming larger towards the east (>50ram). Folded quartz veins are common in the schists, especially in the centre of the profile; they show asymmetric intrafolial folds (Simpson 1986), with sub-vertical axes (N250~176 There are decimetric-scale domains in the psammites with similar folds, and strongly flattened folds in the pelites. Sigmoidal quartz porphyroclasts indicate S-C fabrics, which define a N - S shear zone, the Rio de las Cafias zone (Figs 2, 3). Kinematic indicators define a predominantly dextral-reverse (transpressive) sense of displacement. The described foliation corresponds to $2, with a mean orientation of N31~176 The asymmetric intrafoliar micro-folds are interpreted as resulting from progressive deformation in an essentially non-coaxial stress field, which inverts the sense of stretching compared with the early stages of deformation (Williams et al. 1994). The most constant mineral association in this zone is Bt + Ms + Crd + Mag + Qtz + Plag, with tourmaline, zircon and apatite as accessories. The cordierite porphyroblasts have inclusions of un-orientated biotite (Btl), small laminae of orientated Bt2, and abundant grains of magnetite, quartz and muscovite. They preserve an earlier schistosity (S1) and are surrounded by a lepidoblastic matrix of biotite and muscovite. The two types of biotite described in the previous zone are also recognized here. The paragenesis is Crd + Bt + Ms + Mag + Qtz, and the cordierite was probably produced by the reaction Bt + Qtz = Cdr + Kfs + H20. A n d a l u s i t e + K - f e l d s p a r zone. This zone is restricted to gneiss septa in the central part of the profile (Fig. 3). The growth of K-feldspar and the disappearance of primary muscovite is an important feature. The structures are essentially the same as those that characterize the shear zone of the central sector, e.g. $2 orientated N20"E/66~ The typical mineral association is Bt + And + Kfs + Crd + Mag + Fib + Qtz + Plag, with secondary muscovite, tourmaline and zircon. The gneisses have a granoblastic texture with large poikiloblastic cordierite containing

350

R. J. PANKHURST ET AL.

orientated magnetite and biotite. The feldspar is poikilitic microcline, with myrmekitic borders against plagioclase. Andalusite crystals are 2-3 mm in length, containing magnetite inclusions. Fibrolite is concentrated in folia that anastomose between coarse-grained quartz, plagioclase, K-feldspar, cordierite and andalusite. The paragenesis is Crd + Fib or And + Kfs + Bt and the principal reactions appear to be Ms + Qtz = Fib or And + Kfs + H20, in which andalusite preceded formation of fibrolite. However, textural relations indicate that some andalusite formed after fibrolite during retrograde metamorphism. The scarcity of biotite also suggests Bt + Ms + Qtz + Crd + Kfs + H20. The common presence of the S-C fabric and regional foliation both in the Chepes granodiorite and in the metamorphic rocks suggests that they were generated during a single deformation event, essentially contractional, partly transpressive, which took place over a significant period encompassing the emplacement of the granodiorite.

Anatectic

zone. This has the association Qtz + Plag + Kfs + Crd + Bt + Mag (+Sill + Ms), tourmaline, zircon and apatite. The zone is recognized on the basis of an apparent relationship between the cordierite-rich rocks of migmatitic aspect and the cordierite granites (Tuani Granite) that crop out in the highest part of the sierra (Fig. 3). The granites are characterized by the presence of nodules and euhedral crystals of cordierite with poikilitic nuclei that are intergrown with poikilitic microcline and large laminae of biotite. Myrmekite is developed at the contact between K-feldspar and plagioclase (An20-An~0), which also shows concentric zonation. Sillimanite is very common; muscovite is generally abundant, but late with respect to the rest of the association. The presence of tourmaline, characteristic of the previous zones, is notable. The most probable reaction involves the incongruent melting of biotite: Bt + Sill + Qtz + Plag = Crd + Kfs + melt.

that temperature increased progressively from about 400~ in the biotite zone to about 700~ in the anatectic zone, implying a thermal gradient of 50-100~ 1.

Relationship of metamorphism to deJormation In general, the poikilitic porphyroblasts of the western and central areas possess a relict internal schistosity of S1 type, recorded by inclusion trains of opaque minerals and orientated laminae of biotite and muscovite, at a mean angle of about 20 ~ to the external $2 schistosity. The predominant shape of the porphyroblasts is ellipsoidal, with the main axes parallel to $2. It is important to note that deformation associated with $2 did not rotate the porphyroblasts with respect to geographical coordinates, as indicated by sub-parallel internal S1 trails (cf. Vernon 1988, 1989). The lack of internal crenulation of the S planes and the general straightness of the inclusions trails in cordierite porphyroblasts suggest that cordierite grew during the early stage of development of $2 (Bell & Rubenach 1983; Johnson 1990), implying that mineral growth was essentially syn-deformation. Support for this is seen in the central part of the profile, where some porphyroblasts have a weakly folded S continuous with that of the matrix. Moreover, the mega-porphyroblasts of cordierite include relict micro-folds, indicating that growth of the cordierite outlasted deformation. These observations lead to the overall conclusion that the growth of the M2 porphyroblasts was essentially contemporaneous with the non-coaxial deformation that generated $2 and, additionally, that granodiorite emplacement was contemporaneous with these events.

Granitoids of the Los LlanosUlapes batholith

Conditions of metamorphism The metamorphic zones produced during the M2 event are characterized by (1) the predominance of cordierite, (2) the appearance of andalusite above the stability limit of muscovite and (3) the absence of garnet fi'om the highest temperature zone. This implies a thermal event at low pressures, no greater than 3 kbar (Pattison & Tracy 1991), constraining the depth of emplacement of the granitoids to a maximum of 7-10 kin. The paragenesis of each zone indicates

By far the greatest area of exposed rock in this batholith consists of calc-alkaline granitoids, dominantly metaluminous. The most abundant rock-type is granodiorite, with lesser amounts of biotite monzogranodiorite or monzogranite, but there are also minor amounts of highly peraluruinous granitoids, principally monzogranite with two micas and cordierite. In the Sierra de Los Llanos there are numerous small ellipsoidal bodies of gabbro-diorite and quartz diorite, rich in hornblende; these are the most basic rocks

CENTRAL SIERRAS PAMPEANAS FAMATINIAN ARC

Q A Chepes Granodiorite : Asperezas Granite 'lama Gabbro

Tuani Granite

D

/

V

v

v

v

v

v/

A

P

Fig. 4. Modal QAP diagram for the igneous rocks of the Los Llanos-Ulapes batholith, southern La Rioja. Although a considerable number of analyses are plotted for the basic rocks, this greatly overestimates their relative volume in outcrop (see Fig. 2).

observed in the batholith. Dykes of tonalitediorite in the Sierra de Chepes reach 40m in width but cannot be shown on the map scale of Fig. 2. Five main plutonic units have been distinguished, whose modal compositions are shown in Fig. 4. The high modal quartz content is a characteristic of the Chepes granitoids shared by the granitoids of the Famatina System (Rapela et al. 1992).

(1) Chepes Granodiorite This is the most extensive rock type, frequently possessing a penetrative NW-SE foliation as described in the previous section. In composition it is generally a medium-grained hornblendebiotite granodiorite, with transition to monzogranite and tonalite, at the felsic and mafic ends, respectively (Fig. 4). Occasional small bodies of the tonalitic facies are observed to cut the main granodiorite, indicating complex emplacement of this unit in several pulses. The magmatic mineral assemblage consists of plagioclase (An39-An52), microcline (Or90-Or97), sometimes perthitic microcline, quartz, biotite ({Fe+2/[Fe +2 + Mg]} > 0.44), Mg-hornblende, epidote, allanite, apatite, titanite, magnetite and zircon. There are two main varieties. (i) A biotite-rich variety (with hornblende rare or absent), light grey in colour, of mediumto-coarse grain size (3-6ram). Ellipsoidal enclaves are common and range in composition

351

from metasedimentary inclusions in various stages of assimilation, to dark, fine-grained hornblende-biotite diorite. They are aligned parallel to the foliation. (ii) A dark grey or green variety, rich in mafic minerals, especially hornblende. The latter forms tabular crystals, ~ 5 m m in length. Epidote and sulphides (pyrite or chalcopyrite) are the main accessories. This variety is richer in mafic igneous enclaves, up to 1.5m in length, also orientated parallel to the foliation and flattened perpendicular to it. They may also contain megacrysts of plagioclase feldspar or, less frequently, quartz. Epidote is conspicuous in the Chepes Granodiorite, sometimes with euhedral cores of allanite, and associated with hornblende, biotite, plagioclase, magnetite and titanite. Contacts with hornblende and biotite are euhedral, and the pistacite content is 27-28%, indicating that the epidote is probable primary (Zen & Hammarstrom 1984). However, the depth of emplacement inferred from the metamorphic envelope clearly indicates that the emplacement pressure did not exceed 3 kbar.

(2) Chepes Porphyritic Granodiorite This is a coarse-grained pink biotite granodiorite with rare hornblende, transitional to tonalite and monzogranite. The mineral assemblage is the same as that of the main Chepes Granodiorite and determined compositions are plagioclase (An27-An54), microcline (Or88 Or95), Bt ({Fe+2/[Fe +2 +Mg]} > 0.44) and Mg-Hbl. Its geochemistry extends throughout the range exhibited by the light grey variety of the granodiorite, but is biased towards the more felsic end. It is chiefly distinguished by the presence of megacrysts of alkali feldspar (10 30ram), and allanite and titanite are more conspicuous. Igneous enclaves are relatively scarce, although xenoliths of cordierite hornfels derived from the metasedimentary rocks are found. It has the same penetrative foliation as the granodiorite, and the contact between the two rock types is sharp but welded, implying more or less contemporaneous emplacement.

(3) Asperezas Granite Small (1-5km) round bodies of true granite, normally red or orange in hue, occur throughout the region. They are clearly intrusive into the main granodiorite and/or the metamorphic

R. J. PANKHURST ET AL.

352

(5) Tama Gabbro

rocks, with sharp contacts. The common rock type is monzogranite, transitional to syenogranite, but aplitic varieties occur, mostly as irregular veins. Biotite and muscovite are scarce, with apatite and zircon as the usual accessories.

Small bodies of gabbro, diorite, or quartzdiorite, occur throughout the region, but especially close to the northwestern border of the Sierra de Los Llanos, where they have been exploited as ornamental stone ('black granite'). They frequently show mingling relationships with the surrounding granodiorite, e.g. concentrations of rounded marie igneous enclaves close to the contact. The most homogeneous representative is a dark rock, rich in hornblende, with patchily-zoned plagioclase feldspar (An90An40). The marie minerals are always hydrated, predominantly hornblende, cummingtonite and minor biotite. Cumulate textures are not observed. Accessories include Fe and Ti oxides (up to 3% in the mode), apatite and allanite. The small dykes of tonalite-diorite in the Sierra de Chepes can be included in this unit, as they also show evidence of mingling with the granodiorite, and are rich in hornblende (3560%) and acicular apatite; biotite, epidote and titanite also sometimes occur.

(4) Tuani Granite There are two main facies of peraluminous granitoid that go together under this name. One is a fine-grained orange-brown two-mica granite with conspicuous muscovite (some secondary) and accessory apatite and zircon. The largest body of this facies crops out in the area around Tuani, in the highest part of the sierra (Fig. 2). The second type is a fine-grained brown cordieritic monzogranite, sometimes rich in sillimanite or fibrolite, which is mainly associated with the high-grade metamorphic rocks in the southern part of the Sierra de Chepes, although isolated occurrences also occur between Pacatala and Santa Cruz de la Sierra. This type contains nodules of cordierite, quartz, K-feldspar and biotite, as well as numerous xenoliths of the metamorphic rocks. The accessory minerals are tourmaline (also present in the xenoliths), opaque minerals, secondary muscovite and zircon. Both facies are generally altered. Contacts with the Chepes Granodiorite suggest that they were intruded penecontemporaneously with the latter.

] ] ] I ....

0.2 0.5

....

Many samples (116) were analysed for major and trace element geochemistry. Representative data are given in Table 1, and the broad range of compositions is illustrated in Figs 5-10. The Chepes Granodiorite and Chepes Porphyritic Granodiorite dominate the overall

I . . . . . .

- Asperezas Granite

Peralkaline

A 1.0 o E O~ 0.8 Metalumin~ I ~ < 0.6 2+ C) 0.4

I

I ....

G e o c h e m i s t r y o f the g r a n i t o i d s

~,IL[

IJ

....

] [ III

1 ....

1 ....

] ....

1 ....

0~'~,~0 0

A ChepesGranodiorite

I

....

I I l '

A...~

CaO

Tama Gabbro

0 "~0

Tuani Granite Metamorphic Rocks

-

Peraluminous

9 9

~oo

~

O

0

.0 1.5 2.0 2.5 3.0 AI203/CaO+Na20+K20 (mol)

_~1~

/

qlD

~

~

45 Alkaline

50

Alkaline-calcic

55

60

[.~ [ ~.,_

65

70 75 Calcic

80 SiO 2

Calc-alkaline

Fig. 5. (a) Alumina saturation index (ASI) in the igneous rocks of the Los Llanos-Ulapes batholith, symbols as in Fig. 2. The Tama Gabbro, Chepes Granodiorite and Asperezas granite form a fairly continuous series, mostly metaluminous, but trending into the peraluminous field in the case of the latter unit. The Tuani Granite is strongly peraluminous and is clearly related to the composition of the exposed metasedimentary rocks. (b) Peacock Index (Peacock 1931) for the mainly metaluminous igneous series of (a) (open circles represent CaO v. SiO2, and filled circles total alkalis v. SIO2). The cross-over point falls in the calcic range, as for many subduction-related magma series generally referred to as calc-alkaline.

CENTRAL SIERRAS PAMPEANAS FAMATINIAN ARC

Fig. 6. Ab-An-Or diagram (Barker 1979), with symbols as in Fig. 2. Although the granitoids of the Los Llanos-Ulapes batholith follow a trend typical of active continental margins, the high An content of the more basic rocks is notable and differentiates them from trondhjemitic granitoids of similar age in the Sierras de C6rdoba, farther east (Rapela el al. this volume).

composition of the batholith and are entirely metaluminous (Fig. 5a). Together with the Asperezas granite and the Tama Gabbro, they define a wide-ranging series, from 46% to 79% SiO2 with a Peacock index of 63.5, according to which it should strictly be known as calcic (Fig. 5b). However, the limit of 61 placed by Peacock (1931) on the calc-alkaline series is not generally consistent with present-day usage for the more primitive subduction-related sequences which exceed this limit but are still referred to as calc-alkaline (e.g. Brown 1982). The rocks of the Los Llanos Ulapes batholith have a highly calcic trend in the normative A n - A b - O r diagram (Fig. 6), which distinguishes them clearly from sodic granitoids of the same general age (early Ordovician) in the Sierras de Cdrdoba to the east (Rapela et al. this volume). The granodiorites have medium K20 contents (Fig. 7), and similar general geochemical characteristics to those of batholiths formed at continental margins (e.g. Fig. 10). Their range in SiO2 is 62-72%, ASI 0.85 1.1, total alkalis 3.5-7.0%, [La/Yb]N mostly 3-10, and they have negative Eu-anomalies. Trends show smooth variation (Figs 7, 8) and together with the MREE depletion in the aplites suggests liquidsolid equilibria controlled principally by plagioclase and hornblende (Fig. 9). The igneous enclaves are of tonalitic composition, slightly more basic than the granodiorite (SiO2 58-62%).

353

The Chepes Porphyritic Granodiorite (69-71% SiO2) , which forms a separate mappable unit (Fig. 2), is geochemically similar to the most acidic varieties of the main granodiorite. It does however show enrichment in Ba, Rb and Sr compared with the granodiorite, suggesting effective accumulation of K-feldspar (Dahlquist & Rapela 1997). The Tama Gabbro is the most primitive unit in the Los Llanos Ulapes batholith, with 4654% SiO2. There is a compositional gap of c. 4% SiO2 between the most silicic diorites of this unit and the most basic rocks of the Chepes Granodiorite (enclaves with c. 58%), which, given the extensive sampling, we consider to be real (Figs 7, 8). The gabbro unit is metaluminous (ASI 0.8-1.04, Fig. 5a) and subalkaline (total alkalis 1.5-3.2%). The REE patterns of the most basic samples (up to 48% SiO2; 7-10% MgO) are relatively flat with [La/Yb]N 2.58.0, with no Eu anomalies (Fig. 9); there is a slight depression of MREE, suggesting minor amphibole fractionation. The more acidic varieties of the Tama Gabbro could be derived from these most basic compositions by dominantly plagioclase fractionation, since their REE patterns possess the same slope, but at higher concentrations and with negative Eu anomalies (Fig. 9). The Asperezas Granite, with a SiO2 range of 72-79% and total alkalis of 6.5-9.0%, seems to be continuous with the major element trends exhibited by the granodiorites (Fig. 7). It has slightly higher ASI (0.96 1.12) but this rarely exceeds a value of 1.1, the limit for I-type granitoids of the Lachlan Fold Belt of Australia (Chappell & White 1992). Whereas some trace elements (Rb, Ga, Y) also seem to lie on extensions of the granodiorite trends, others (Ba, Sr, Th) show an abrupt decrease with increasing SiO2 (Fig. 8). Rapid decrease of Sr and Ba in granitic series is commonly attributed to removal of K-feldspar and biotite (e.g. Chappell 1996; Rapela & Pankhurst 1996). The REE patterns mostly show [La/Yb]N = 2.5-6.5, with a few values up to 13, and negative Eu-anomalies. The aplites on the other hand have relatively flat REE patterns ([La/Yb]N= 0.5-1.5) and depletion in total REE compared with the granodiorites, probably due to fractionation of a REE-rich accessory phase such as allanite, which is present in the Chepes Granodiorite and conspicuous in the Chepes Porphyritic Granodiorite. The Tuani Granite is locally associated with high-grade metasediments (Fig. 2), and its geochemistry and petrogenesis are clearly distinct from those of the other granitoids. It is depleted in CaO and Na20, and enriched in

354

R. J. P A N K H U R S T ' '1""1''"1""1""1''"1'"'

'"'1'"

0 20 -,, II~

'"1''"1

A1203

9

9

m

5

~ 1 7 6 1, 7 6 ,,,,I,,,

10 , | 9

TiO2 1.5

=m z~ z~

10

,I,,,,I,,,,I,,,,I,,,,I,,,,

'"1''"1'"'1'"'1'"'1''"1"

;

~c~

15

,],

ET AL.

1.0

e~o +

b,-

,,I,,,,I

,,,l,,,,l,,J~l,,,+-~-~rl,,,,l,,

CaO -

FeO*

9

8

9

Na20 +

~ o ~:~ 9

MgO

- - I-

+

10

9

, ~ + - + ~, 9

12

9 -,_

9 m~Pi

++

o ="

-

4

z~z~

o

o

~

high-K

00%

++

+

K20

.,,~.;-~+

~

9

.~.,~*

A ~

9

P205

~.

,

o.2 t

+mediurn-K

A "Ao4m~=

+

o 4..,

9~ 50

0,

,ow.~ ~ ,ow., 55

60

65

70

SiO 2

75

80

85

50

55

60

65

70

75

80

85

SiO 2

Fig. 7. Harker variation diagrams for major elements in the granitoids of the Los Llanos-Ulapes batholith (symbols as in Fig. 5a). For most elements, the rocks of the calcic metaluminous series show coherent trends that can be explained by continuous crystal fractionation of the principal rock-forming minerals. P20s is an exception, showing a trend with negative correlation for the granodiorites and granites; this occurs in low-K series, or calcalkaline series poor in P, where apatite crystallization is restricted to the most evolved rocks (Green & Watson 1982). The fields shown in the K20 plot are those of Le Maitre et al. (1989).

C E N T R A L SIERRAS P A M P E A N A S F A M A T I N I A N A R C

;;oo o

30

20

355

Cs

Ga 0

0 0o

0

12

Zk

++

9

+~ +

%.^+% 10

9

,"

A

~",~~+

# ~ mm

~ 0

,,I,,,,I

.... I .... ( .... I,,,,l,,,,I

U

0

o

-

++

o

o

i,

oo

9

o

"

~

_

A ~+,~+ z~ +

~

o

%

+

" - - ~++~+~ +

150

~..~

A~A~'~Z'~x ~ ,

200

+~=++ zx.~',-# +

9

o

o

~%

Rb

+++

o

0000

9 +

oo

%*

.... I,,,

+

o

o

~ m + ++

Z:b~:~#~ll

+

100

o

=_~

o o

50

iw30

+'t-

Th

,,

-

A 0

o ~.~ OOAo "~ ~ ~

10

9

A

A

0% 0

20

Sr

AA

200

A

A

+ + ++++ .-'o+

i-11, ~

A#

-".~-- % ~ o _

A ZS 7 7 .

o

~o

Z~

o

o

100

"

"

,,

I.... l,,,,l~,i,l,,,, Y ,,,,I .... ll,i, ,,Lil,,,,li,,,l .... I...._L. . . . . . +

A 0

,~,

Ba

A 0

0

AA 9

0 0

-H+ + + -~4~-

40

o 0 oO~O

-

."

800

ix

600

9 n~

-

~

400

o

I I

0

20

z~'~ a

zh + + ~

~

AA

d ~

,,,I .... I .... I .... I,,,,I .... I.... liii,lill 60

65

70

Si02

75

80

85

o ~'+~

r 55

50

I.... I,,, +

Y

+ +.+ :-I-+

50

o

o

60

150

o

,r 9 ~

+~

,',,,l .... ~,,,,l,,,,I,,,,• 50

55

60

.... ~ ....

65

70

75

0

§ l ....

80

0

200

~,,,

85

SiO2

Fig. 8. Harker variation diagrams for trace elements in the granitoids of the Los Llanos-Ulapes batholith (symbols as in Fig. 5a). Trends for the LIL elements (Sr, Ba and Cs) have an inflexion at about 68-73% SiO2, which can be explained by the initial crystallization of a major phase, such as K-feldspar. Rb shows a continuous positive trend, emphasising its incompatible character. HFS elements (Th, U and Y) have inflexions at 72-75% SiO2, indicating changes in the crystallization of the principle accessory phases (allanite, titanite, apatite and zircon) (Bea 1996). Ga shows negative correlation, as does A1203.

356

R. J. P A N K H U R S T

Rock/Chondrites I

lOO -

I

I

I I

I

I

I

I

Tama G a b b r o

10

ET AL.

I

I

I

9 CAL-71 i LC-1 0 AGN-27 TA-25

I

I

I

I

I

Chepes

I

46,3% 48,4% 51.8% 54.2%

.....

T~" ~

I

I

I

I

I

i

I

I

I

I

I

Granodiorite ]A GUA-254 61.9% 0 NAC-256 66.3% .... .......

.........

ApLites

",,, / ~

& RLC-236 66.9% 9 PUL-244 68.8% , I ') PAS-272 69.5%

I-- SAN-225 71.0% - -

....... .............._ "l--

i-

--!t-!

[::', NAC-258 77.0% I 1

I

I

I

I

I

Asperezas

100 --

[

I

I

I

Granite

I

I

t

I

I

I

I

I

I

I

I

I

I

I

I

I

I

I

I

L I

~ AGU-217 74.5% ,J 9 GUA-249 75.2%

, -203

757O/o1

J

10 ....

~"............~ Aplites ~ ~,~ ]

I

La Ce

I

I

Nd

I

L

I

~ AGU-221 78.9% GUA-252 77.9% ELE-208 75.0% i

I

I

SmEu Gd Tb Dy

I

I

Er

I

L

]

Yb Lu

I

I

9 TUA-247 65.9% 0 PAS-277 66.2%

Metamorphicr o c k s [] EKE-211 58.9%

2 TUA-246 67.4% QUI-111 68.1%

9 RLC-231 60.0%

I

La Ce

I

I

Nd

I

L

I

I

I

I

SmEu Gd Tb Dy

I

Er

I

I

I

Yb Lu

Fig. 9. Chondrite-normalized rare earth element patterns of the granitoids of the Los Llanos Ulapes batholith K 2 0 , A1203, TiO2, FeO and MgO compared

to the calc-alkaline series (Fig. 7). It has a restricted range of SiO2 (66-69%) and very high ASI (1.2-1.8), in accordance with its strongly peraluminous mineralogy. Notably, in the case of the feldspar-forming elements (A1203, CaO, N a 2 0 and K20) there is an apparent continuity with the pelitic rocks of the metamorphic basement (Fig. 7). The REE patterns have [La/Yb]N of 6.0-7.8), with negative Euanomalies, and are indistinguishable from those of the metamorphic rocks (Fig. 9).

1000

-

s

n-CO,O

Rb 100

10

G

-

Geochronology

Metasedimentary rocks Samples from the profile described above (Rio de las Cafias) and from close to Malanzfin (Fig. 2) were analysed for Rb-Sr systematics (Table 2, Fig. l la). Scatter considerably in excess of that allowed by the measurements is indicated by an MSWD of 59, resulting in a high enhanced error on the age calculated from the errorchron, 513 :~ 31 Ma. Nevertheless, this

IllJl

I

1

I ,,I,,,I

I

10

~ l l,~,,r

Yb+Ta

loo

Fig. 10. Rb v. Yb + Ta plot (Pearce et al. 1984) for the granitoids of the Los Llanos-Ulapes batholith. The metaluminous series fails mainly in the volcanic arc granite field (VAG), with the Asperezas granite, the most evolved group, extending into the syn-collisional (syn-COLG) and within-plate granite (WPG) fields. The less siliceous Tuan~ granite plots along the boundaries of these fields.

CENTRAL

SIERRAS PAMPEANAS

FAMATINIAN

ARC

357

Table 2. Rb-Sr data for igneous and metamorphic rocks from the Sierras de Los Llanos de La Rioja Sample

Rb

Sr

87Rb/~6Sr

%or

87Sr/'86Sr

%ff

(87Sr/86Sr)o

134 119 154 136 137 126 150 128 225 190 149 142 227 117 119 143 120 153 152 134

2.3103 3.0364 2.4149 3.3071 4.0491 2.4296 1.6761 2.3568 1.1047 1.3340 2.1841 2.1427 1.0878 3.3212 3.2658 2.5565 3.7260 2.0191 2.6292 2.8607

0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5

0.725834 0.730764 0.725608 0.732267 0.733866 0.725477 0.720431 0.724304 0.716542 0.717966 0.724221 0.723888 0.714906 0.731472 0.731262 0.726122 0.733078 0.721663 0.727654 0.730224

0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010 0.010

0.70970 0.70956 0.70875 0.70918 0.70559 0.70851 0.70873 0.70785 0.70883 0.70865 0.70897 0.70893 0.70731 0.70828 0.70846 0.70827 0.70706 0.70757 0.70930 0.71025

149 165 141 132

1.9090 0.9807 1.4195 1.9713

0.5 0.5 0.5 0.5

0.722820 0.715261 0.718365 0.721644

0.010 0.010 0.010 0.010

0.70949 0.70841 0.70845 0.70788

189 35

2.0552 18.6122

0.5 0.8

0.721501 0.833130

0.010 0.020

0.70715

137 146 151 157

0.5654 0.7696 0.4130 0.7116

0.5 0.7 1.2 0.7

0.714418 0.715974 0.711322 0.714120

0.010 0.010 0.010 0.010

0.71047 0.71060 0.70844 0.70915

47 49 38 84 12 18 18 71 78 83 60 58 19 44 51 55 52 82 20 93 107 13 12

11.8409 9.0421 12.5470 3.6941 59.2022 48.4441 34.1086 7.4316 6.7230 5.1907 9.2253 9.7289 25.0541 13.3064 12.7450 8.6080 9.9182 5.7049 29.8466 4.4938 3.9306 38.8680 41.1772

0.5 0.5 0.7 0.5 1.6 1.3 1.3 0.5 0.5 0.5 0.5 0.5 1.2 0.6 0.5 0.5 0.5 0.5 1.2 0.5 0.5 1.6 1.7

0.787969 0.769663 0.791917 0.733428 1.082020 1.030332 0.940016 0.758540 0.754081 0.743574 0.769808 0.771663 0.875299 0.795840 0.792600 0.766088 0.774968 0.747939 0.905595 0.739214 0.735847 0.970196 0.986201

0.015 0.010 0.015 0.010 0.025 0.025 0.025 0.010 0.010 0.010 0.010 0.015 0.020 0.020 0.020 0.010 0.010 0.010 0.020 0.010 0.010 0.025 0.025

Chepes Granodiorite (490Ma) ILI-222 ILI-223 ILI-224 SAN-225 RLC-233 RLC-234 RLC-236 RLC-243 GUA-253 GUA-254 NAC-255 NAC-256 MIS-259 PAS-261 PAS-262 PAS-269 PAS-272 PAS-271 OLT-283 OLT-285

107 124 128 155 190 106 87 103 86 87 111 104 85 133 132 125 152 106 136 131

EncNves (490Ma) OLT-284 RLC-237 RLC-238 RLC-242

97 56 69 90

Porphyritic Granodiorite (490 Ma) PUL-244 PUL-245

134 224

Tama Gabbro (490Ma) LLA-286 LLA-287 LLA-289 LLA-290

27 39 22 39

Asperezas Granite (470Ma) ELE-203 ELE-204 ELE-205 ELE-206 ELE-207 ELE-208 RLC-239 RLC-240 RLC-241 NAC-257 NAC-258 GUA-249 GUA-250 GUA-251 GUA-252 OLT-279 OLT-280 OLT-282 OLT-281 PAS-263 PAS-264 PAS-266 PAS-267

190 152 166 107 230 293 200 182 180 147 188 195 160 202 225 162 175 159 203 142 143 174 165

0.70911 0.70869

0.70878 0.70906 0.70733 0.70540 0.70651

0.70726 0.70845 0.70855 0.70974 0.70912 0.70953

R, J. P A N K H U R S T E T AL.

358 Table 2, (continued) Sample

Rb

Sr

Asperezas Granite (470Ma) PAS-268 215 PAS-273 197 PAS-273b 188 PAS-274 179 PAS-275 199 PAS-270 182 Tuani Granite (490Ma) TUA-246 125 TUA-247 184 PAS-277 166 PAS-276 140

t~7Rb/86Sr

%(7

STSr/86Sr

%(7

0.5 0.5 0.5 0.8 0.5 1.8

0.788542 0.772721 0.760710 0.810007 1.104810 0.756368

0.015 0.010 0.010 0.020 0.025 0.010

2.6386 4.5402 5.3785 3.4184

0.5 0.5 0.5 0.5

0.733208 0.741379 0.750411 0.736706

0.010 0.010 0.010 0.010

0.71554 0.70968 0.71286 0.71284

2.6348 1.5652 7.9946 2.9011 6.2696 6.2322 4.3992 3.9549 1.7732 4.5991

0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5 0.5

0.733139 0.724118 0.771700 0.734411 0.759223 0.757063 0.747598 0.740287 0.725709 0.745087

0.010 0.010 0.015 0.010 0.010 0.010 0.010 0.010 0.010 0.010

0.71474 0.71319 0.71588 0.71283 0.71415 0.71070 0.71545 0.71355 0.71688 0.71367

(continued) 53 12.0675 61 9.5022 72 7.6618 35 15.2687 10 57.8467 74 7.2114 139 119 90 120

Metamorphw Rocks (at 490Ma) RLC-226 99 109 RLC-227 112 207 RLC-228 206 75 RLC-230 82 82 RLC-231 217 101 RLC-232 106 49 RLC-229 168 112 ELE-209 161 118 ELE-210 65 106 ELE-21I 134 84

(87Sr/S6Sr)0

0.70909 0.70940

Rb and Sr concentrations in ppm, determined by XRF at BG.S. Initial S7Sr/S6Sr calculated at time given in parentheses, except where 2(7 error exceeds 0.2%. sent case, since the age overlaps with that determined for the Chepes granodiorite, and some o f the samples come from the zones clearly affected by contact m e t a m o r p h i s m , an unambiguous interpretation of the e r r o r c h r o n age is not

result is consistent with an e r r o r c h r o n result of 523 + 18 M a obtained for the Los Tflneles phyllites in the Sierras de C 6 r d o b a (Rapela et al. this volume) and equally consistent with a C a m b r i a n regional m e t a m o r p h i s m . In the pre-

(a)

Malanzdn

(b) 87Sr/ eeSr

0.80- 87Sr/86Sr Schists and Phyllites 0.75 0.78

o El Elefante 9 Rfo de las Cafias ~

_

Age 513 + 31 Ma I.R. 0.713 _+0.001 MSWD = 59 Enhanced Errors

0.74

~ 0.72 S

0.74

~ Tuani Granite o Granodiorites Enclaves : Tonalites

0.73

[] Isochron p o i n t s ~

J

0.76

Chepes Granodiorite (all areas)

0.72

=/./i/ ~

_

0.7089:1:0.0002

~

f

479 + 6 M_a

-~'"

[MSWD = 2.6] (Reference isochron

0.71 J

for 9 of 24 points)

87Rb/ S6sr I

I

2

I

I

4

I

I

6

I

I

8

~

I

10

87Rb/ 8SSr I

,

0

:

,

:

:

~

1

:

:

',

~

1

2

:

~

~

'

;

3

~

:

:

:

',

4

:

:

:

'.

',

:

:

5

Fig. 11. Rb-Sr isochron plots. (a) Low-grade metasedimentary rocks from the Rio de las Cafias and E1 Elefante 1 km north of Malanz',in. The errorchron is comparable to that obtained for similar rocks in the Sierras de C6rdoba (Rapela et al. this volume), both in age and initial S7Sr/86Sr. The result is thought to reflect the same Cambrian metamorphic event, in the present case disturbed by contact metamorphism during emplacement of the igneous rocks. (b) Chepes Granodiorite: data points for 19 granodiorite and tonalite samples and four of enclosed dioritic enclaves scattered widely about the reference isochron shown, which is defined only when more than half the data are excluded.

C E N T R A L SIERRAS P A M P E A N A S F A M A T I N I A N A R C

359

Table 3. Sm-Nd data for igneous and metamorphic rocks, Sierras de Los Llanos de La Rioja

147Sm/144Nd 143Nd/144Nd ( 143"Nd/ 144Nd)t

cNdt

Tchur

TDM*

Chepes Granodiorite (490 Ma) RLC-243 5.491 28.327 GUA-253 5.038 20.922 NAC-255 5 . 4 4 1 29.902 OLT-283 6.923 31.688 PAS-272 5.456 27.524 MIS-259 8.845 40.006 SAN-225 5.169 25.569

0.1172 0.1455 0.1100 0.1321 0.1198 0.1336 0.1222

0.512123 0.512214 0.512078 0.512133 0.512119 0.512200 0.512110

0.511747 0.511747 0.511725 0.511709 0.511734 0.511771 0.511718

-5.1 -5.1 -5.5 -5.9 -5.4 -4.6 -5.7

991 1268 988 1195 1033 1063 1084

1608 1608 1639 1662 1626 1574 1650

EncOres (490Ma) RLC-242 4.807 OLT-284 6.489

0.1354 0.1238

0.512192 0.512106

0.511757 0.511709

-4.9 -5.9

1113 1116

1593 1662

Porphyritk Granodiorite (490Ma) PUL-244 7.331 36.535 0.1213

0.512183

0.511794

-4.2

924

1541

Tama Gabbro (490Ma) LLA-290 3.755 16.918

0.1342

0.512138

0.511707

-5.9

1222

1664

Asperezas Granite (470 Ma) ELE-206 6.114 25.959 OLT-279 4.920 19.068 OLT-282 5.171 24.199 RLC-241 4.848 25.304 GUA-252 1.787 5.448 NAC-257 3.845 18.391 PAS-273 3.950 17.066

0.1424 0.1560 0.1292 0.1158 0.1983 0.1264 0.1399

0.512173 0.512221 0.512142 0.512085 0.512370 0.512159 0.512141

0.511735 0.511741 0.511744 0.511728 0.511760 0.511770 0.511710

-5.9 -5.7 -5.7 -6.0 -5.4 -5.2 -6.3

1309 1566 1124 1045 (-ve) 1043 1338

1648 1640 1635 1657 1613 1598 1682

Tuani Granite (490 Ma) TUA-246 7.836 39.230 PAS-276 6.480 33.508 PAS-261 6.898 32.626

0.1207 0.1169 0.1278

0.512051 0.512067 0.512113

0.511663 0.511692 0.511703

-6.7 -6.2 -6.0

1561 1482 1577

1725 1686 1671

PhyHites (at 490Ma) ELE-209 8.347 43.361 RLC-228 7.617 39.326

0.1164 0.1171

0.512007 0.512011

0.511633 0.511635

-7.3 -7.3

1559 1564

1767 1764

Sample

Sm

Nd

21.461 31.686

Sm, Nd concentrations in ppm, by MSID at NIGL. *calculated according to DePaolo et al. (1991).

possible. This overlap does not apply however to the Los Tfineles case, where it clearly dates a M i d d l e - L a t e C a m b r i a n event, most probably the P a m p e a n regional m e t a m o r p h i s m . The 'initial' 87Sr/86Sr ratios of both groups at this time are also relatively high (0.713-0.714) and, assuming that deposition was not very m u c h older than the m e t a m o r p h i s m , this suggests p r o v e n a n c e of the sedimentary material from m a t u r e continental sources. S m - N d m o d e l ages for crustal residence of 1760-1770 M a (Table 3) indicate a possible age for this ultimate source as old as late Palaeoproterozoic.

Chepes Granodiorite R b - S r dating of the main Chepes granodiorite was not successful (Fig. 1 lb). Despite a reasonable range in R b - S r ratios, the data do not c o n f o r m to an isochron, p r e s u m a b l y reflecting

the complex intrusive and genetic history of this widespread and variable rock type. The fit is not i m p r o v e d if the data are treated as subsets representing samples from single localities, implying either that that the m a g m a s were isotopically heterogeneous or that the R b - S r systems in the rocks have been partially reset. The general trend of the data array corresponds to a b o u t 4 8 0 M a , with an initial 87Sr/86Sr of a b o u t 0.710. If the most discordant data points are progressively excluded from the fit, for no other reason than their statistical scatter, an isochron is achieved with 9 o f the 19 data points, yielding an 'age' of 479 + 10 Ma. There is no distinction between the enclaves and their host granites in terms of initial Sr isotopic composition (Table 2). U - P b zircon dating was carried out on granodiorite and a mafic enclave within it, from the central zone of Rio de las Cafias (Table 4). The data are plotted in Fig. 12a,

360

R. J. P A N K H U R S T E T AL.

,_o ,,o c,i ,.o ea e.l

F. r

%

l'~

-H -H +1 -H +l

~,

~r ~

r O0 xO

,..o

xO s c,"l t ~ O~ 0% tCb P'-~ 0 xO I'~ ~ 0

"0

C . ~ .O . O . O . O

t~

e,i

i

,.o

o ,,..:, oo

,.o ,,o cq

-~

%

~

cq

,.o

v~

,..o xt ,.o ~q

9

E~

~

',o

9 ~"

~.~

tr

.~

9

._~

' 5 o .

,.~

~_

~-

e',q

%/

~

~

~

,-..~ ' ~ - .-.-~ ~

,---,

~t'q

r

~.~

~

~.

CENTRAL SIERRAS PAMPEANAS FAMATINIAN ARC

(a)

361

(b) 0.070 9

9

/

(")

0.082

-

~Z)

M.'ge

0.080 9~

a_

Granodiorite .------~" (equantgrains) , , ' /

M

~

RLC-236

- _~

:/"0400

,__

- .

m.

Ma,iceoc,ave

~

~

n

o

d

i

o

r

i

uongn e e a , e s ) / q ~ 0 0

t

e

/

i

Zircon

n=

20

I

[ 4 9 7 +_ 6 M a

0.9

t 0.065

9 600

0.078 Granodiori~e" " - - ~ _ . . ~ 0 . 1 3 ^ Chep..es..

', , ,

' I

_x

T

.Upper ~ Intercept 22(?O-Ma_

0.060

~

I I~P~'tI~LX460

,

,,o,

,

0.055

0.076 //,,," "'~ 0.58

1

0.60

~

0.62 207 p b / 235U

CordieriteGranite

0.64

0.66

10

11

12 13 238 U / 208p b

14

0.050

Fig. 12. U-Pb zircon systematics for the Chepes Granodiorite. (a) conventional abraded fractions from the granodiorite and an included enclave plot along a discordia which cuts Concordia at 486 + 4 Ma (2or). The lower intercept, poorly defined, is at c. 100 Ma. (b) SHRIMP data for the same granodiorite sample are essentially concordant (hr error bars are shown) and give a mean m6pb/23SU age of 496 + 7 Ma (2or). These two ages are combined to give an overall crystallization age of 490 + 5 Ma. which shows that four fractions, including one from the enclave, are concordant at c. 490 Ma, albeit two with large errors due to high common Pb corrections. Two other fractions are collinear on a discordia line, one apparently having suffered Pb-loss, the other lying above Concordia, which may signify U-loss. The discordia line ( M S W D = 2 . 0 ) intersects Concordia at 4 8 6 + 4 M a , which is taken as the age of crystallization of the granodiorite. The mafic enclave shows no memory of any earlier history. The same granodiorite was also dated by U - P b S H R I M P analysis of the rims and tips of uniformly zoned, euhedral zircons, yielding a mean 2~ age of 4 9 6 + 7 M a (Table 5, Fig. 12b). It is possible that this age is slightly overestimated due to the inclusion of data influenced by older inheritance, since there is some scatter to the left of the main cluster in Fig. 5; there was no optical or cathodo-luminescence evidence of inheritance in the analysed crystals.

Other granitoid rocks Independent dating of the Porphyritic Granodiorite was not carried out, partly because the sample sets were too small, although the wholerock and contiguous aplite listed in Table 2 (PUL-244, PUL-245) give a two-point age of 473 + 9 M a . This is indistinguishable from the R b - S r errorchron and conventional U - P b ages for the main granodiorite, but is outside the error range of the S H R I M P age.

For the Tuani Granite even fewer samples were available and these, being clearly S-type, anatectic granites, show variable inherited compositions that preclude simple dating. Thus, the R b - S r data points scatter above the granodiorite isochrons (Fig. 1 lb), and the single analysed zircon fraction plots in a highly discordant position in the U - P b concordia diagram (Fig. 12a), despite abrasion prior to analysis, suggesting inheritance from a Palaeoproterozoic component. R b - S r data for the Asperezas granite were assessed by sub-dividing into three localised geographical groups (Fig. 13). Two of these groups, from around Malanzfin and the southern part of the Sierra de Los Llanos respectively, give ages close to 470Ma, although only the former can be regarded as a true isochron (MSWD values are 2.8 and 5.0, respectively). The initial 87Sr/S6Sr ratios of both are close to 0.7090. The third group, from the eastern part of the Sierra de Chepes, is even more scattered (MSWD = 5.5), and yields both a younger age (452 • 11 Ma) and a higher initial 8VSrff6Sr ratio (0.7105 + 0.0013).

Crystallization history and comparison with Famatinian arc The most reliable ages for crystallization of the Chepes Granodiorite should be the two U - P b zircon results (conventional and SHRIMP). Together they give a weighted mean age of

362

R. J. PANKHURST E T AL.

Table 5. Summary of ion microprobe U-Pb results for zircons from sample RLC-236 Spots

U

Th

Th/U

(ppm) (ppm) 1.1 1.2 2.1 3.1 4.1 5.1 6.1 7.1 8.1 9.1 10.1 11.1 12.1 13.1 14.1 15.1 16.1 17.1 18.1 19.1

117 333 288 125 393 264 348 161 316 310 426 270 310 258 438 306 337 285 295 508

82 327 203 128 332 150 213 150 187 258 249 208 270 185 334 166 235 186 188 374

Pb*

2~176

2~

l~r

(ppm) 0.70 0.98 0.71 1.02 0.85 0.57 0.61 0.93 0.59 0.83 0.58 0.77 0.87 0.72 0.76 0.54 0.70 0.65 0.64 0.74

18 54 43 20 62 39 50 25 46 50 60 40 47 38 71 44 52 41 41 78

0.000007 nd 0.000121 0.000098 0.000128 0.000443 0.000023 0.000183 0.000158 0.000078 0.000085 0.000168 0.000126 0.000145 0.000078 0.000140 0.000515 0.000112 0.000050 0.000118

0.0825 0.0805 0.0792 0.0797 0.0816 0.0803 0.0794 0.0803 0.0793 0.0846 0.0779 0.0781 0.0774 0.0775 0.0856 0.0797 0.0832 0.0776 0.0753 0.0813

0.0019 0.0026 0.0024 0.0027 0.0129 0.0035 0.0019 0.0035 0.0017 0.0068 0.0044 0.0033 0.0026 0.0038 0.0088 0.0023 0.0013 0.0013 0.0044 0.0014

Age(Ma)

2~

lo"

511 499 491 494 506 498 493 498 492 524 484 485 480 48l 529 494 515 482 468 504

11 16 14 16 77 21 12 21 10 40 26 20 16 23 53 13 8 8 26 8

nd, 2~ detected. Correction for common Pb made on the basis of extrapolation to concordia along a mixing line with common Pb, following Tera & Wasserburg (1972), as outlined in Compston et al. (1992).

4894-3Ma, although the fact that they only overlap within 30- errors suggests that these errors may be underestimated. As noted above, it is also possible that the SHRIMP age is slightly biased to an older age by cryptic inheritance not distinguished analytically. A more conservative figure rounded off to 4 9 0 + 5 M a is therefore preferred. Thus emplacement of the Chepes Granodiorite, in this area at least, took place in earliest Ordovician times (Tremadoc according to the time scale of Gradstein & Ogg 1996). On the basis of coherent field relations and geochemistry, a similar age may be assumed for the same granodiorite unit in other parts of the batholith, and for the Tama Gabbro with which they exhibit magma mingling. The Rb-Sr systems for both the main Chepes granodiorite and the porphyritic variety appear to have been partially reset at about 475-480Ma, which would also help to explain their heterogeneity with respect to Sr isotopes. Interpretation of the data for the Asperezas Granite is more difficult in the absence of U - P b results for these rocks. The 450Ma errorchron age for one group is most probably reset, judging by its higher initial S7Sr/86Sr ratio. The other two may represent a separate episode of granitoid emplacement in mid-Ordovician times (470Ma would be Arenig/Llanvirn), and the Asperezas Granite is the one unit of the Los

Llanos-Ulapes batholith for which field relations would permit a significantly younger age. On the other hand, they could have been emplaced at around 490 Ma immediately after the granodiorite, and been partially reset in midto Late Ordovician times. The close chemical coherence in the Harker diagrams suggests a genetic relationship between all rock groups (except the Tuani Granite), with a common process of fractional crystallization of the major igneous phases from a series of comparable (cogenetic) magmas. In either case, the R b - S r ages are comparable to those that have been obtained from the Famatina System to the northwest: 456 + 14Ma for Cerro Toro (Saavedra et al. 1992), 451-t-28 Ma and 4 5 7 + 18 Ma for granodiorite and granite respectively from the Sierra de Paganzo (Saal 1993; Saal et al. 1996). Loske & Miller (1996) also report U - P b zircon data for the Nu15orco granite in the central Sierra de Famatina that are discordant but suggest a crystallization age of close to 460 Ma. The most conservative interpretation of the new data is that the bulk of the g a b b r o granodiorite magmatism of the Los LlanosUlapes batholith occurred at c. 490Ma, some 25-30 Ma earlier than previously suspected, but that similar subduction-related magmatism recurred at 470 Ma and may even have persisted

CENTRAL SIERRAS PAMPEANAS FAMATINIAN ARC

(a)

1.20. f 1.10

=u)

1.00

~~

Malanzfin G r a n i t e s

]

F El Elefante ~

0.90

Age 469 + 5 Ma I.R. 0.7089 + 0.0004 MSWD = 2.8

0.80

87Rb/86Sr

:::::::::::::::::::::::::::::::::::::: 0

10

20

30

40

50

f

0.90,

o9 r

~/o

oo o Guasamayo [

I" Nacate 'L~ Olta

0.85. 0.80,

~ / /

I ]

- - - "

_/J_

_

~Age452+11Ma I.R. 0.7105 + 0.0013

~

MSWD = 5.6 Enhanced Errors

0.75, ~ 0.701 . . . . 0

: .... 5

(C)

1.20-

1.10

60

(b) E a s t e r n C h e p e s G r a n i t e s

0.95

: .... 10

.... 15

: .... 20

:

87Rb/86Sr : .... : 25 30

Pacatala - Santa Cruz

-~

1.00 ~ 0.90

Age 471 +_8 Ma I.R. 0.7091 + 0.0008 MSWD 510 Enhanced Errors

~ J

=

~

0"80

87Rb/86Sr 0.7(

:

0

I,

10

i

I

20

,

I

30

~

I

40

,

I

50

,

I

'

60

I

70

Fig. 13. Rb-Sr for members of the Asperezas Granite unit, grouped roughly according to geographical location. The western (a) and northeastern (e) groups give consistent ages of c. 470 Ma, but the central group at least may have suffered subsequent re-homogenization.

until 450Ma, both here and in the Sierra de Famatina. The tectonic significance of the magmatic history in the two areas is further considered below. Subsequent Devonian S-type granite magmatism is known in both areas.

Sm-Nd isotope geochemistry The analyses presented in Tables 2 and 3 indicate that the high initial S7Sr/a6Sr values of the granitoids is matched by relatively low eNdt

363

(-4.6 to -6.7). Only the Tuani Granite, with initial 87Sr/86Sr ratios of 0.710-0.716 and cNdt of < - 6 , is close to being identifiable as a direct crustal melt of the local country rock, perhaps with some mixing with the isotopically more primitive calc-alkaline suite. From the field and geochemical evidence we interpret these highly peralkaline granites as largely resulting from local anatexis in the roof zone of the batholith. The source of the Asperezas Granite, with a mean 87Sr/86Sr470 of 0.7086 and gNd470 of -5.6, is indistinguishable from that of the granodiorites and diorites, which have mean values (at 490 Ma) of 0.7084 and -5.3, respectively. (If the age of the Asperezas granite is assumed also to be 490 Ma, the mean eNdt is still -5.5, but the initial S7Sr/S6Sr becomes much less radiogenic, at 0.7066). Perhaps the most remarkable aspect of the isotope compositions is that the mafic and intermediate rocks of the calc-alkaline suite, even including the Tama Gabbro, share the same crustal signature (Tables 2 and 3). There is no evidence for massive in situ contamination during emplacement in the upper crust, so this signature must reflect the composition of the middle or lower crust, regardless of whether this was incorporated as a result of magma generation in these regions, or to heavy contamination of the Rb-Sr and Sm-Nd systems during ascent. In either case, emplacement through mature continental crust is an inevitable conclusion. Sm-Nd crustal residence ages were calculated using the multistage model of DePaolo et al. (1991), which gives more realistic ages for crustderived magmas than the usual model for direct derivation from depleted mantle. It is seen that all the major units have a common well-defined model age of 1550-1750Ma (Table 3), which is largely an indication of their generally uniform eNdt values. This is essentially the same range as that exhibited by the Cambrian plutonic rocks of the Sierras de Cdrdoba (Rapela et al. this volume) and strongly suggests that the underlying continental crust is of the same age and composition in both regions (Fig. 14). On the other hand, the Famatinian (Inner Cordilleran) granitoids of C6rdoba have lower Nd model ages of 1200Ma (Rapela et al. this volume), signifying a younger and/or more primitive source. Apart from the latter, the crustal model ages of the central and Eastern Sierras Pampeanas are clearly older than the Meso/Neoproterozoic ages derived for the hidden basement of the Precordillera and the Western Sierras Pampeanas (Fig. 14), which are taken as indicative of a Grenvillian allochthonous terrane.

364

R. J. PANKHURST E T A L .

Sierras de C6rdoba

10

[--] G1 Igneous Suite 1

Metamorphic Basement

Sierras de Los Llanos de La Rioja

Tuan[ Granite

15

10

(+21 Asperezas

[~ ] 9 Precordlllera[ (xenoliths) & / Sierra de Pie I de Palo .J

~+I I, Granite

~ Tama ~ Gabbro

5

II[i]

- mr~ ~:.~ Chepes Granodiorite o o '...0

O o co

o o o

O o 04

o o ~

o O (.0

o o cO

o O o

o o 04

Tdm* Ma Fig. 14. Sm-Nd model ages. Histograms of the multistage model ages from Table 3, for the time of differentiation from depleted mantle of the crustal reservoir from which the magmas of the Los LlanosUlapes batholith were derived. The strong concentration around 1600-1700 Ma for all rock types coincides with a similar peak for the Sierras de C6rdoba. The latter includes both metamorphic basement and pre- to syn-Pampean igneous suites which were in part anatectic (Rapela et al. this volume). It is concluded that the underlying crustal structure of both areas is identical. It contrasts with the younger (Grenvillian) crust of the Precordillera terrane, indicated by the Sm-Nd model ages of xenoliths in the Precordillera itself (data from Kay et al. 1996) and the 1021 + 12Ma gneisses of the Sierra de Pie de Palo (Pankhurst & Rapela this volume; authors' unpublished Nd data).

Conclusions and geotectonic interpretation of the Famatinian arc The geochronological, mineralogical and microstructural data presented here indicate that the

predominant granodioritic stage of magmatism in the Los Llanos-Ulapes batholith occurred at 4 9 0 • in earliest Ordovician times, synchronously with the imposition of the regional $2 schistosity in a contractional tectonic regime. Evidence for magma mingling between the Chepes Granodiorite and the Tama Gabbro require that the latter are also essentially of the same age. R b - S r data for the monzogranitic stage (Asperezas granite) suggests a younger age of emplacement at c. 470450Ma, i.e. mid- to Late Ordovician times, although this needs confirmation by more extensive U - P b dating. The timing and nature of this calc-alkaline magmatism is relevant to the evolution and significance of the Famatinian arc to the north, as presently understood. The geological parallels between the two regions are very strong. In the Sierra de Famatina, the metamorphic basement to the east also consists of low-grade metasediments comparable to those described here, as well as higher-grade gneisses of pelitic origin which have given a R b - S r age of 523 + 47 Ma and an initial 87Sr/86Sr ratio of 0.715 (Saal 1993; Saal et al. 1996). The metagreywackes (Negro Peinado Formation) have been interpreted as derived from a mature igneous-cratonic source (Clemens & Miller 1996). It is likely that most of these pre-arc Pampean sequences could be considered to be broadly equivalent to the Cambrian metasediments of the western Sierras de Cdrdoba (Rapela et al. this volume) and the Puncoviscana Formation of the foreland. As noted above, they were apparently all laid down on similar Palaeo/Mesoproterozoic basement rocks prior to the Pampean orogeny (see Pankhurst & Rapela this volume). It may also be noted that that the metamorphic conditions, and the relationships of metamorphism to the deformation and the emplacement of the granitoids, are similar to those described for the southern part of the Famatina System. In the Sierra de Paganzo, the southernmost part of the Famatina System and closest to the present study area, Saal (1993) deduced high-T, low-P (3-3.5 kbar) metamorphism in the host rocks of the granitoids. As a whole, the magmatic rocks of the Sierra de Famatina are more varied than those of the Los Llanos-Ulapes batholith. Deposition of the post-Pampean siliciclastic sequences of the former, spanning the Tremadoc to Early Llanvirn interval, are predominantly marine shales, calcareous sandstones and cherts but they are interbedded with volcanic rocks. The latter comprise a bimodal basalt-rhyolite association, interpreted as initially an island arc

CENTRAL SIERRAS PAMPEANAS FAMATINIAN ARC sequence, passing upwards into more acid rocks as the arc became integrated with the continental margin (Toselli et al. 1996). However, even the basalts have initial 87Sr/86Sr ratios of 0.7070.7095 (Mannheim & Miller 1996), indicating a crustal component as in the igneous rocks of the Los Llanos-Ulapes batholith. The plutonic rocks of the Famatina System are broadly similar, in some cases almost identical, to those described here. In the Sierra de Paganzo (Saal 1993), tonalites and granodiorites with mafic enclaves predominate, emplaced at shallow depth concomitantly with ductile deformation and the development of a major shear zone. These were subsequently intruded by monzogranites interpreted as co-magmatic. Farther north, the central zones of the Famatina System are dominated by the granitoids of Nufiorco-Sanagasta, Cerro Torre and Palm,n, which also exhibit the same range of calcic or calc-alkaline compositions (see individual papers in Acefiolaza et al. 1996). In some areas, basic rocks (gabbros and norites) are relatively common, and in others there are large volumes of cordieritic granitoids that resemble those of the Sierras de Cdrdoba. The available geochronological data for these rocks, mostly R b - S r ages of 450-460 Ma, were noted above, but the possibility that some of these might reflect postemplacement re-homogenization is raised by the new data for the Los Llanos-Ulapes batholith. Lork and Bahlburg (1993) presented midOrdovician U - P b monazite ages of 467 476 Ma for granitoids intruded into Puncoviscana Formation and its equivalents in the Puna, north of the Famatina System. Our main conclusions from these comparisons and the data presented in this paper are: 9

9

9

the Los Llanos-Ulapes batholith is a southern continuation of the Ordovician calcalkaline, subduction-related, magmatic arc generally known as the Famatinian arc; the rocks were emplaced into a metasedimentary rock sequence equivalent to the Puncoviscana Formation, s.l., which had been deposited on westward facing continental slope and had been metamorphosed to very low grade during the Pampean orogeny in Early to Mid-Cambrian times (see Rapela et al. this volume); the main emplacement of voluminous metaluminous Chepes granodiorites and the Tama gabbro occurred at 490 + 5 Ma, considerably earlier than previously supposed, towards the end of regional compressive deformation and shearing in the metasedimentary rocks ($2 fabric). Locally in the roof

9

9

9

365

zone, temperatures were high enough to result in anatexis to form the S-type Tuani Granite unit; this was followed by post-deformational emplacement of smaller monzogranitic plutons. Rb Sr data suggest that this intrusive phase took place at c. 470Ma, and may have persisted until 450Ma, although geochemical continuity with the granodiorites and the possibility of resetting of R b - S r errorchrons might mean that this unit (Asperezas Granite) was also Early Ordovician in age; all the magmas show strong isotopic evidence for emplacement through mature continental crust, with Nd model ages of 1600-1700Ma. In this area at least, therefore, the Famatinian magmatic arc is clearly a continental one, not an island arc; the lower crust of this region appears to be identical to, and presumably therefore continuous with, that of the Sierras de COrdoba, the supposed eastern foreland of Gondwana in Ordovician times.

Although considerable work remains to be done in elucidating the history and origin of magmatism in the Famatinian arc, we believe that considerable modification is required to the idea that this was formed entirely in response to a m i d - O r d o v i c i a n ocean closure related to the approach of the allochthonous Precordillera terrane. This work was carried out as part of contract CI1CT92-0088" from the Commission of the European Communities and was also supported by grant PIP No. 4148 from the Consejo Nacional de Investigaciones Cientificas y T6cnicas de la Rep~blica Argentina (CONICET). It is presented as a contribution to IGCP Projects 345 (Andean Lithospheric Evolution) and 376 (Laurentia-Gondwana Connections), and constitutes NIGL Publication No. 239. We thank M. Brown and C. F. Miller for helpful reviews.

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BEA, F. 1996. Residence of REE, Y, Th & U in granites and crustal protoliths; Implications for the chemistry of crustal melts. Journal of Petrology, 37, 521-552. BELL, T. H. & RUBENACH, M. J. 1983. Sequential porphyroblast growth and crenulation cleavage development during progressive deformation. Tectonophysics, 92, 171-194. BROWN, C. G. 1982. Calc-alkaline intrusive rocks: their diversity, evolution, and relation to volcanic arcs. In. THORPE, R. S. (ed.) An&sites." Orogenic Andesites and Related Rocks. John Wiley, New York, 437-464. C~APPELL, B. W. 1996. Compositional variation within granite suites of the Lachlan Fold Belt: its causes and implications for the physical state of granite magma. Transactions of the Royal Society of Edinburgh: Earth Sciences, 87, 159-170. & WHITE, A. J. R. 1992. I- and S-Type granites in the Lachlan Fold Belt. Transactions of the Royal Society of Edinburgh: Earth Sciences, 83, 1-26. CAMmOS, R. 1979. Sierras Pampeanas Noroccidentales. Salta, Tucumfin, Catamarca, La Rioja y San Juan. H Simposio de Geologla Regional Argentina. Academia Nacional de C6rdoba, C6rdoba, I, 225-291. CLEMENS, K. A. & MILLER, H. 1996. Sedimentologia, proveniencia y posici6n geotect6nica de las sedimentitas del Precfimbrico y Paleozoico inferior del Sistema de Famatina. In: ACE~qOLAZA et al. (eds) q.v., 31-50. COMPSTON, W., WILLIAMS, I. S., KIRSCHVINK, J. L., ZHANG ZICHAO & MA GUOGAN 1992. Zircon U-Pb ages from the early Cambrian time-scale. Journal of the Geological SocieO,, London, 149, 171-184. DAHLQUIST, J. A. & BALDO, E. G. 1996. Metamorfismo y deformaci6n Famatinianos en la Sierra de Chepes, La Rioja, Argentina. XIII Congreso Geolagico Argentino y III Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, V, 393409. & RAPELA, C. W. 1997. Los granitoides ordovicicos de la Sierra de Chepes, Batolito de Los Llanos-Ulapes, Sierras Pampeanas. XIV Reuniao de, Geologia do Oeste Peninsular, Vila Real, Portugal, Comunicagoes, 31-36. DALLA SALDA, L., CINGOLANI,C. & VARELA, R. 1992. Early Paleozoic orogenic belt in southwestern South America: result of Laurentia-Gondwana collision? Geology, 20, 617-620. DEPAOLO, D. J., LINN, A. M. & SCHUBERT, G. 1991. The continental crustal age distribution; methods of determining mantle separation ages from Sm-Nd isotopic data and application to the Southwestern United States. Journal of Geophysical Research, B96, 2071-2088. GRADSTEIN, F. M. & OGG, J. 1996. A Phanerozoic time scale. Episodes, 19, 3-5. GREEN, T. H. & WATSON, E. B. 1982. Crystallization of apatite in natural magmas under high pressure, hydrous conditions, with particular reference to 'orogenic' rock series. Contributions to Mineralogy and Petrology, 79, 96-105.

JOHNSON, S. E. 1990. Lack of porphyroblast rotation in the Otago schist, New Zealand: implications for crenulation cleavage development, folding and deformation partitioning. Journal of Metamorphic Geology, 8, 13-30. KRETZ, R. 1983. Symbols for Rock-Forming Minerals. American Mineralogist, 68, 277-279. LLAMB[AS, E. J., SATO, A. M., ORTIZ SUAREZ, A. & PROZZI, C. 1998. The granitoids of the Sierra de San Luis. This volume. LE MAITRE, R. W., BATEMAN,P., DUBEK,a., KELLER, J., LAMEYRE, J., LE BAS, M. J., SABINE, M. A., SCHMID, i . , SORENSEN, H., STRECKEISEN, A., WOOLLEY,A. R. & ZANETTIN, B. (eds) 1989. A Classification of Igneous Rocks and Glossary of Terms. Recommendations of the International Union of Geological Sciences Subcommission on the Systematics of Igneous Rocks. Blackwell Science, Oxford. LORK, A. & BAHLBURG,H. 1993. Precise U-Pb ages of monazites from the Faja Eruptiva de La Puna Oriental, NW Argentina. XII Congreso Geol6gico Argentino y H Congreso de Exploraci6n de Hidrocarburos, Mendoza, Actas, IV, 1-6. LOSKE, W. & MILLER, H. 1996. Sistemfitica U-Pb de circones del granito de lqufiorco-Sanagasta. In: ACEIqOLAZA et al. (eds) q.v., 221 227. LUDWIG, K. 1980. Calculation of uncertainties in U-Pb isotope data. Earth and Planetary Science Letters, 46, 212-220. MANNHEIM, R. & MILLER, H. 1996. Las rocas volcfinicas y subvolc'anicas eopaleozoicas del Sistema de Famatina. In: ACENOLAZAet al. (eds) q.v., 159-186. PANKHURST, R. J. & RAPELA, C. W. 1998. Introduction. This volume. , SAAVEDRA,J., BALDO, E., DALQUIST,J. & PASCUA, I. 1996. Sierras de Los Llanos, Malanzfin & Chepes: Ordovician I- & S-Type granitic magmatism in the Famatinian Orogen. XIII Congreso Geol6gico Argentino y II1 Congreso de Exploraci6n de Hidrocarburos, Buenos Aires, Actas, V, 415. PASCUA, I., SAAVEDRA, J. & BALDO, E. 1996. Metamorfismo y magmatismo del Paleozoico inferior en la zona central de las Sierras Pampeanas (ST de Gondwana). IV Congreso Geol6gico de Espafia, Alcal~ de Henares, 20, 1672-1675. PATERSON, S. R., VERNON, R. H. & TOBISCH, O. T. 1989. A review of criteria for the identification of magmatic & tectonic foliations in granitoid. Journal of Structural Geology, l l , 349 363. PATTISON, D. R. M. & TRACY, R. J. 1991. Phase equilibria and thermobarometry of metapelites. In: KERRICK, D. M. (ed.) Contact Metamorphism. Mineralogical Society of America Reviews in Mineralogy, 26, 105-206. PEACOCK, M. A. 1931. Classification of igneous rocks series. Journal of Geology, 39, 54-67. PEARCE, J. A., HARRIS, N. B. W. & T1NDLE, A. G. 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of Petrology, 25, 956-983.

CENTRAL SIERRAS PAMPEANAS FAMATINIAN ARC RAMOS, V. A., JORDAN, T. E., ALLMEND1NGER,R. W., MPODOZIS, C., KAY, S. M., CORTI~S, J. M. & PALMA, M. 1986. Paleozoic terranes of the central Argentine-Chilean Andes. Tectonics, 5, 855-880. - - 1 9 9 6 . Los klippes y ventanas tect6nicas prefindicas de la Sierra de Pie de Palo (San Juan): edad e implicaciones tect6nicas. XIII Congreso Geol6gico Argentino y Ill Congreso de Explorao cian de Hidrocarburos, Buenos Aires, Actas, V, 37%391. RAPELA, C. W. & PANKHURST, R. J. 1996. Monzonite suites: the innermost Cordilleran plutonism of Patagonia. Transactions of the Royal Society of Edinburgh: Earth Sciences, 87, 193-203. , COIRA, B., TOSELLI,A. J. & SAAVEDRA,J. 1992. The Lower Paleozoic magmatism of southwestern Gondwana and the evolution of the Famatinian orogen. International Geology Review, 34, 1081-1142. , PANKHURST, R. J., CASQUET, C., BAEDO, E., SAAVEDRA, J., GALINDO, C. & FANNING, C. M. 1998. The Pampean Orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de Cdrdoba. This volume. SAAL, A. 1993. El basamento cristalino de la Sierra de Paganzo, provincia de La Rioja, Argentina. PhD Thesis, Facultad de Ciencias Exactas Fisicas y Naturales, Universidad Nacional de Cdrdoba. , TOSELLI, A. J. & RossI DE TOSELLI, J. N. 1996. Granitoides y rocas b~,sicas de la Sierra de Paganzo. In: ACEIqOLAZAet al. (eds) q.v.199 209. SAAVEDRA, J., PEELITERO-PASCUAL, E., RossI DE TOSELLI, J. N. & TOSELLI, A. J. 1992. Magmatic evolution of the Cerro Torre granite: a complex Ordovician pluton of northwestern Argentina. Journal of South American Earth Sciences, 5, 21-32.

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Index Note: Page numbers in bold type refer to Tables, those in italic type refer to illustrations, diagrams etc Ab-An-Or diagram, Sierras Pampeanas 353 Acadian orogeny 77, 156, 228 Acado-Baltic Province 68 Acanthotoechia 64, 73 accretionary prisms 137 accretionary stages 3, 13, 70, 75, 226 Achala batholith 184, 185, 338 Achalian orogeny 261,278 actinolite-tremolite 286 active margin Famatina System 72, 222, 285 major element evidence 138 advective transfer 223 Aegiromena 65, 67, 70 AFM diagram, Famatina plutons 288 Afro-South American brachiopod Realm 77 aggradational environment 52 agnathans 65, 70 Ahtiella Zone 64 Alabama-Oklahoma transform fault 18 Alcaparrosa Formation 24 pillow-lavas 95 algae 42, 90 allochthons, Precordillera 1, 11, 27 Alsate Shale 93 Altai Hinggan mobile belt 63 alumina saturation index diagram 352 Amazonia 73, 77, 211 American Realm 62 amphibole barometry, Sierra de Fiambalfi 303 amphibolite-facies metamorphism Sierra de Fiambalfi 303-305, 313, 321 Sierra de Pie de Palo 159, 160, 177 amphibolites MORB signatures 212 Sierra de Pie de Palo 160, 168, 172 Sierra de San Luis 236 spider diagram 187 anatexis E1 Pil6n 204, 207, 208, 210 Famatina System 286 Olta 275 Sierra de Chepes 350 Sierras de Cdrdoba 228 Ancasti 297 Anchoramena 65 andalusite 241,242, 243 andalusite zone, Sierra de Chepes 349 Andean cycle 2, 3 Andes, geological map 252 andesite 92, 93, 96, 128 Angaco belt 146 Anglesey 72 Annamitella 73 Anti-Atlas 66 Antirhynchonella fauna 76 apatite 130 Apheorthis 66 aplogranite 198 Aporthophyla 64 Appalachian basin 231 Appalachian-Caledonide orogen 72

Appalachian-Ouachita orogenic belt 226 Appalachians Chepultepec interval 42 correlation with 5, 14, 57, 77 deformation 156 internal basement massifs 15 mobile belts 219 passive margin 16, 74 rift-drift transition 213 Taconic orogeny 225 227 Appohimchi Subprovince 77 apron, volcaniclastic 128 Ar/Ar dating analytical methods 267, 298 Chilenia collision 177 micas 273-275 Sierra de Fiambalfi 308, 316 Sierras Pampeanas 147-149, 150-151,152-153, 154, 155, 222, 261,268 spectra 274 Arbuckle Limestone 64 arc accretion 228 arc splitting 178 arc system, Ordovician basin 127 Archaeorthis Zone 63 Arenig brachiopods Famatina and Central Andean basin 62 Precordillera 61 Arequipa-Antofalla craton and Amazonian craton 211 collision 76, 220, 251 distinctness 6 isotopic ages 73 Occidentalia terrane 5 supercontinent break-up 212 Argentina geological map 298 geological provinces 86 Precordillera 1 Arica embayment 73,220 Artejia mega-continent 70 Asaphiscus 68 ash beds intercalation 20, 27 Precordillera 107 ash transport, direction 120 Asperezas Granite 351-352, 353, 362, 364 aulacogens 15, 211 Avalonia terrane 73, 74, 76, 229 see also Carolina terrane Avalonian faunas 49 B. naris Zone 63, 66 Back Bay fauna 76 back-arc environments Brasiliano orogeny 262 Famatina System 72 Lau basin 170, 171 Pie de Palo complex 146, 175, 178 Precordillera basement 144 Puna 127 Quebrada Guayaupa-Lima 173

370 back-arc environments (continued) Sierras de C6rdoba 214 spreading centres 172, 178 back-arc protoliths, analyses 164-165 Bahamas 20 Balcarce Formation 111 Baltic Province 49, 57, 63, 64, 66 Baltica, bentonites 86 Baltica palaeocontinent 68 barometry, GRIPS 307, 308 basalts discrimination diagrams 287 lavas 128, 172 MORB-type 46, 48, 160 basement rocks Sierra de Pie de Palo 87, 143, 159 Sierra de San Luis 236 basic/ultrabasic belts 229 basin closure 21 basin subsidence 128 bedforms 39 Belen schists 211 Be#ella 66 Bemberg 241,327, 329 K-bentonites 20, 21 ages 110 Baltica 86 comparative chart 120 comparative distribution 121 dating 108 Famatinean volcanic arc 121,229 Fort Pefia 96 geochemistry 107, 115 localities 108, 109 Marathon 93, 96 mineralogy 111 North America 121 Ordovician 7, 75 Solitario 95 transects 110 whole-rock analyses 118-119 see also meta-bentonites Big Bend 89 bioherms, distribution 98 biostromes, San Juan Formation 43 biotite zone, Sierra de Chepes 349 bioturbation 43 Birmingham graben 18, 19, 41 Bishop Tuff 113 bivalves 66 black shales 75, 90 Black Warrior Basin 18 block-faulting 20, 25, 92, 98, 100, 253 Blue Ridge 18, 226, 227, 228, 229, 231 Bonnia-Olenellus chron 51 boudinage 240, 247 brachiopods affinity variation 65 assemblage zones 61 dispersion 59, 70 Famatina and Central Andean basin 62 faunas 57, 59 Gondwana 49 Hirnantian 65

INDEX Laurentia 63 Ordovician 14 Precordillera 61, 63 Brazil, cycles 3 Brazilian orogeny 73, 261 break-up stages 3 break-up unconformities 49, 52 breccias, hydroclastic 128 Brevard fault 226 Caborca 231 Cacheuta sub-basin 146 Cadomian terrane 229 calc-alkaline plutons 225, 226, 277, 290, 318 calc-silicate rocks 313 calcarenites 92 calcretes 40, 90 Caledonian orogeny 76, 77 Calingasta 108 Calmayo 198 Cambrian, faunas 11 Camerella 64 Campo Chico 77 Campylorthis 65 Cafiani 228 Cancafiiri Formation 67 Carlo del Oeste 77 Carlo Grande Formation 77 Capilla del Monte granite 184 Caradocian brachiopods 65 Precordillera and Central Andean basin 63 Carap6 shear zone 185 carbonate facies, warm-water 71 carbonate rim, North America 20 carbonate sequences Cambrian & Ordovician 14, 15, 35 environmental reconstruction 23 facies types 15 Ordovician 20 and rifting models 19, 219 carbonate shoal complex 39 Carolina terrane 74, 75, 76, 226, 227, 229, 231 cataclasis 160, 167 catch-up system 41 cathodo-luminescence images 269 Caucete Group age of 160, 177 contact 146 Cuyania 88 quartzites 155, 169 Sierra de Pie de Palo 98 Celtic Province, brachiopods 57, 59, 63, 64, 72 Central Andean basin, biogeography 66-67 Central Batholithic Belt 3, 184 Central Newfoundland volcanic sequence 73, 78 Cerro Aspero batholith 184, 338 Cerro Barboza 146, 159, 160 Cerro Blanco granites 290 Cerro La Chilca 109 Cerro La Silla 91, 109 Cerros Largos 242, 243, 332, 333 Cerro Negro 313, 314, 315 Cerro Pelado 41, 90, 195

INDEX Cerro Potrerillo 109 Cerro San Jorge 143, 146-147 Cerro Toro granite 288 Cerro Totora Formation 17, 18, 38, 41, 51 Cerro Totora graben 19 Cerro Valdivia 145, 155, 159, 160, 166 Cerro Viejo 95, 109, 111, 114 Chafiic phase 3, 229 channel systems 128 Charlotte belt 226, 227, 231 chemical index of alteration 131, 140 Chepes Granodiorite 351,353, 362 emplacement 365 U-Pb dates 360 Chepes Porphyritic Granodiorite 351,353 Chepultepec interval 42 chert beds 93 Chile trench 1 Chilenia terrane 5, 52, 155, 177, 220, 278 chromite 163, 166, 178 chromitites 183 chronostratigraphic chart 16 Clarkeia fauna 50, 52, 70 clastic rocks palaeocurrents 285 Puna 129 clastic wedge, oceanic crust in 15 clasts basement 45 volcanic 92, 93 Clavohamulus hintzei 63 clay mineralogy bentonites 111 Fort Pefia 97 Cliftonia oxoplecoides 65 climate, Ordovician 20 climate indicators, Precordillera migration 14 clinopyroxenes, compositional plot 115 clitambonitids 66 collision dates, Precordillera terrane 6-7, 12, 75 collision and exchange model 70 Colohuincul complex 224, 225, 228 compositional layering 305, 310 compressional deformation 50 Conglomerado Cafiada Honda 245 conglomerates chert-pebble 48, 49 continental margin 46 mineral compositions 320 Sierras Pampeanas 319 wedged 24 conjugate margins 17, 19, 59, 68, 213 Conlara Metamorphic Complex 236 conodont faunas 35, 43, 63 in brachiopod zone calibration 61 contact metamorphism 306, 311,328 continental arc systems 7, 131, 137, 139, 343, 365 continental bridges 14, 15 continental collision 1, 211,212, 214, 219 magma generation 338 continental lithosphere 14 continental margins, Atlantic-type 47, 52 continental platform, carbonate 13

continental reconstruction Cambrian 227 Early Ordovician 26 continental rise 49 continental shelf, Laurentia 11 continental slope 49, 91, 96 contourites 46 cooling, mid-Palaeozoic 313-314, 316-317 cooling rates 312-313 coquinas 39 cordierite zone, Sierra de Chepes 349 cordieritite 197, 198, 207, 208 Cordillera de la Costa 228 Cordillera Frontal 220 Cordillera Oriental 66, 76 Cordylodus cf. angulatus 43 correlation chart, stratigraphic units 60 Cow Head Group 63 crenulation 184, 245, 250, 306 Crepicephalus Zone 62 cross bedding 42 cross-stratification, herringbone 39, 40 Crossiskenidium 66, 73 crust isotopic signatures 207 partial melting 292 crustal extension Cambrian 14, 51, 52 Ordovician 99 supersequence C 48 crustal recycling 140 crustal shortening 120 crustal subsidence 44 crustal thickening 210, 212, 222, 224, 227, 228 Sierra de San Luis 337 Cruz de Carla 243, 332, 333 Cruz del Eje complex 275 Cruziana 111 Cumbres Calchaquies 222 cumulates Cerro Barboza 175, 178 Cerro Valdivia 168, 175, 178 Mogote Corralitos 173 Pie de Palo complex 160, 167, 171 Quebrada del Gato 175, 178 Sierra de Fiambalfi 312 Cuparius 63, 69 Cushamen unit 225 Cuyania terrane 3, 6, 36 basement 44, 50, 220 carbonate platform 97 collision 53, 155 components 87 composite nature 147 drifting 253 eastern margin 235, 253 extent of 85 faunal exchange 68 Grenvillean age 147 as 'Llanoria' 87 plate rotation 100 rifting 177, 178 Cuyania-Gondwana suture 156, 253 Cuyo basin 146

371

372

INDEX

dacite 92, 93, 128, 131, 174, 286 Dagger Flat Sandstone 92, 93, 98 Dalmanella testudinaria 65 Davidsville Group 66 debris flows 85 Los Sombreros Formation 21 volcaniclastic 93 deformation, compressive 241,242, 245, 253 Degamella 66 dehydration melting 336 Deicke beds 121 depositional assemblages 38 descending slab, hydration of 292 Destombesium 67 detachment models 19 detrital quartz 90 Devil's River uplift 92, 94, 96 diamictites 52, 70 diatexite 187 diorite 328, 350 disruption, post-docking 25 distal slope sediments 92 Dob's Linn 108 dolomitization 40 Dominica 96 Don Braulio Formation 47, 52, 60, 65 Don Polo Formation 46 Douglas Lake 121 Drabovia 65, 70 Drabovinella 67 Dreisbachian faunas 62 drowning carbonate bank 20, 22, 52, 75 Cerro Pelado 90 San Juan Formation 43 ductile deformations, Sierra de Pie de Palo 147, 154, 155, 169, 177 ductile shear 184, 185, 212 Dunnage zone 76 dyke complex 170, 172 dykes aplitic 315, 335 cross-cutting 241 La Escalerilla granite 328, 330 lamprophyre 290 pegmatite 241,242, 250, 315 Phyllite Group 236 rhyolite 329 Sierra de San Luis 328 tonalitic 185 E-MORB 172 Eastern Basement Complex 240, 243, 245-250 Eastern Cordillera 3, 7 Eastern Sierras Pampeanas 3, 5, 7, 220 Ediacaran fauna 74 Ehmania 68 El Diquecito group 186, 187, 192, 197, 210 E1 Estero Member 41 E1 Gigante 228 E1 Hornito 335 E1 Horno Formation 76 E1 Jaguelito terrane 225, 228

E1 Jaguelito-Nahuel Niyeu metapelites 225 E1 Morro pluton 335, 336 E1 Paso Group 98 E1 Pefi6n Grande 335 E1 Pil6n complex age 204, 222, 276 concordia plot 208 emplacement 207 metamorphism 185 migmatites 193, 197 monzogranite 183 El Quemada 178 El Realito 327, 328, 330 E1 Relincho Formation 90 E1 Salado 335 El Telarillo 335 Empozada Formation 25, 91 end-member deposits 21 end-run hypothesis 14, 25 endemism Arenig 72 Caradoc 65 Eorhipidomella 67 Eostropheodonta hirnantensis 65 erosional unconformities, high-angle 47 eruptive plumes 112 Espinal Formation 283 Estonia 63 Euorthisina 66 European Realm 65 eustatic control 44, 51 evaporites 17, 18, 38, 88, 90 extension, lithospheric 21, 24, 27 facies associations 38 facies variation Marathon/Solitario basin 97 marine volcano 94 Faja Eruptiva de la Puna 73 Faja milonitica a Arenilla 247 Famatina System 3, 5, 7, 21, 72 age of 95 basement 283-285 cross-sections 291 evolution 292, 293 geological map 284 granite geochemistry 289 petrogenesis 290 plutonic rocks 288-290, 365 review 283 subduction 213, 214, 251,277, 292 volcanism 286-288 western margin 253 Famatinean cycle 223, 235, 240 243, 251,283 Famatinean magmatic arc 7, 59, 70, 72, 213 active margin 228 ages 277 bentonites 12I continental arc 343 crystallization age 361-363 early stage 182 in Iapetus Ocean 219 interpretation 364-365

INDEX Famatinian orogeny 2, 3, 25, 122, 184 accretion 277 continental collision 219, 261 deformation 223 evolution 227-229 granitoids 327-332 map 221 metamorphism 204 Ocloyic phase 327 Pampean cycle 222 plutons 209 Taconian equivalence 228-231 Famatinolithus 66 Famatinorthis 66, 73, 74 fan deltas 50 fault overstepping 25 fault scarps 46, 47, 48, 99 faulting, syn-depositional 93, 97 faunal compositions 46, 49 faunal divergence 71 faunal exchange 70 faunal migration 67 faunas, Cambrian 11, 14 feldspars, compositional plot 115 ferromanganese crusts 172 Ffynnonia 66, 73 fibrolite 190, 350 fifth-order cycles 40 Finnmarkian orogeny 74 flat-slab subduction 183 flat-slab zone, Nazca plate 1, 5 flexure, lithospheric 20 flood basalts, Paranfi-Etendeka 3 flooding surfaces 39, 43 Florida basement, Gondwana 229 flute marks 128 flysch 50, 52, 236 foliation Sierra de Chepes 348 Sierra de Fiambalfi 301,307, 311 Sierras de C6rdoba 184 Yaconian 229 foreland, Gondwana 1 foreland basin Puna 76 Punta Negra 146 Fort Cassin Formation 63 Fort Pefia Formation 93, 96 fossil hash 39 fossils, La Laja Formation 38 framework minerals, turbidites 128 framework modes, turbidites 130 Franklin Mountains 94, 98 funeral ship scenario 59, 219 gabbro Sierra de Fiambal~t 308 Tama 352 gabbro-norite 311, 312, 313, 315, 318 Gander terrane 74, 75, 76 garnet, textures 190, 301-302, 303, 305 Gasparillo 241,245, 327 gastropods 42, 90 General Alvear sub-basin 146

373

geochemistry bentonites 107, 115 Lower Turbidite System 134-135 Puna turbidites 131 Sierras de C6rdoba 188, 189 Upper Turbidite System 136 Volcanosedimentary Succession 132-133 geochronology, Sierras Pampeanas 356-361 geodynamic scenario, Sierras de C6rdoba 212 geotraverse 181,343 glaciation, Gondwana 14, 27, 47 glauconite 93 global reconstruction 90 Glossopleura chron 51 gneisses analyses 166 Pie de Palo complex 173 174 Quebrado del Gato 174 Sierra de San Luis 327 see also orthogneiss, paragneiss Gogoella 66 Gondwana docking of Precordillera 11, 155 faunas 50 Florida basement 229 foreland 1, 3, 7 glaciation 14, 27, 47 Pampean collision 212 pre-Andean margin 77 proto-Pacific margin 1, 68 rotation 57 western margin 222, 278 see also Laurentia-Gondwana Gondwanan Iapetus 57 Gondwanian cycle 2, 3 graben filling 24 graben systems 19, 23, 49 graded beds 41 grainstones 20, 42, 43 Grampian orogeny 74 Grand Pitch Formation 74 granites hypersolvus 338 Sierra de Fiambalfi 315 317 within-plate 3 granitization 228 granitoids classification 195 evolution 337 Famatinian 223, 327-332 geochemistry 224, 289, 352-356 Los Llanos-Ulapes batholith 350-352 Mamil Choique 225 metaluminous 195-197, 202, 227 peraluminous 7, 183, 197-198, 204, 207, 213, 227, 276, 352 post-orogenic 184, 223, 334-336 Puncoviscana Formation 227 Sierra de San Luis 325, 330 Sierras de C6rdoba 183 Sierras Pampeanas 343 syn-orogenic 332--334 tectonic fields 331,333 Western Sierras Pampeanas 122, 223

374 granodiorite Famatina System 290 La Playa 207 Quebrada de Gato 174 Rio Salado 225 Sierra de Chepes 348 Sierra de San Luis 328 Sierras de C6rdoba 183, 195 syn-kinematic 334 granulite facies 183, 210, 240, 312 graptolites 93, 97, 99, 285 gravity flows 98 greenschist facies Sierra de Fiambal/l 304, 307, 313 Sierra de Pie de Palo 159 Sierra de San Luis 249, 250, 253 Grenville orogeny 73 Grenvillian Cuyania 147 Famatina basement 285 granites 209 Marathon/Solitario 92 Occidentalia 222 Pb signatures 145 Precordillera 12 San Rafael 44 Sierra de Pie de Palo 220 U-Pb dates 87, 161 Western Sierras Pampeanas 3, 5, 6, 85 greywackes, Rhenohercynian 131 Gualcamayo Formation 90, 108, 109, 111, 120 Guamanes Shear Zone 266, 275, 278 Guandacol bentonites 108, 122 carbonates 88 Formation 23, 38 hardgrounds 41 lithoclasts 90 sediment thickness 47 sedimentary environment 43 tectonic phase 228 unconformity 41 Gfiiraldes 198, 207, 208 Guyana craton 76, 77

Hallograptus etheridgei 97 hardgrounds, Guandacot 41 Harker variation diagrams 354, 355 harzburgites 183 Hawke Bay event 41, 44, 51,226 Hayesville fault 227 heat-flow, lateral 49 heavy minerals 50, 52, 131 Hebridean shield 73 Hesperonomia 64 Heterorthis 67 Hf-Th-Ta plot 171 HFS elements 139, 184 high-Na plutons 195, 198, 207, 209, 213 highstands 39, 43 Hirnantia fauna 70, 76 H. sagittifera 65, 67 hornblende, Ar/Ar ages 149, 150-151, 154 horst and graben structures 22-23, 48, 99

INDEX

Huacoella Zone 63, 64 Hungioides 66 hydrothermal activity, spreading ridges 172 L victoriae lunatus zone 122 I-type source rocks 277 Iapetus Ocean 6, 7, 12, 14 closure 76, 227 evolution 77, 219 margins 57 obduction 100 opening 15, 73, 74, 225, 226 palaeogeography 25 width 108 Iapetus suture 226, 229 ice-rafting 52 Idiostrophia 64 Illaenopsis 66 Illaenus 73 immobile elements 159, 172, 177 in-situ partial melting 312 incompatible elements, enrichment in 184 Incorthis 66, 67 Indian Bay terrane 78 intraclasts 42 Inversella (Reinversella) 64 Irish Northwestern terrane 73, 75 Iruyic event 229 Iruyica unconformity 222-223 island arcs 20, 127, 144, 146 Isograptus caduceus 93 isotopic data, Sierras de C6rdoba 185 Jfichal 43, 44, 49, 108, 111,120 Jan Mayen Ridge 26 Juan XXIII granite 183 Jujuyapis 66 K-Ar ages San Jorge 147 Sierra de San Luis 251,333, 335 Sierras de C6rdoba 184, 187 K/Ca ratios 151 kaolinite 111 karstification 94 King Mountains 231 Kinnekulle beds 120 kinzigites 195, 208 Kosov Province 65 kyanite 240, 250 La Aguadita Formation 285, 293 La Bolsa 238 La Calera group 186, 187 La Cantera Formation 21, 47, 60, 94 La Carolina 245, 250 geological map 239 La Chilca block 23 La Ci6naga 332, 333 La Cienaguita 335 La Cruz limestones 41 La Cumbre-Ascochinga 185, 195, 202

INDEX La Escalerilla granite 236 age 278, 332 chemistry 332 dykes 328, 330 emplacement 242, 243, 330 margin 245, 249 La Estanzuela 186, 187, 197, 210 La Flecha Formation 40, 90 La Florida 241,242, 250 La Fronda 198, 207, 208 La Invernada Formation 21 La Laja Formation 38, 88 La Melada complex 247, 250 La Melada-La Gruta 238 La Pampa 85, 143 La Playa 198, 207 La Puntilla 305, 306, 311, 313, 318, 321 La Represa 332, 333 La Rioja Malanz/m Formation 262 thrust belt 12, 111 La Silla Formation 41-42, 63, 90 La Tapera 332 La Totora pluton 335 La/Sc v. Ti/Zr plot 139 La/Ta ratios 160, 166 Lachlan Fold Belt 353 lagoonal deposits 39 laminites 90 lamprophyre dykes 290 Las Aguaditas Formation 22, 23, 44, 60, 91 Las Aguilas Group 236, 249, 264, 277 Las Chacras batholith 332, 334, 336, 337 Las Chacritas section 44 Las Pirquitas thrust 147, 155, 160, 169 Las Planchadas Formation 122, 287, 289 Las Plantas Formation 21, 60 Las Trancas 223 Las Vacas Formation 21 Las Verbenas 241,243, 327, 330 Lau Basin 159, 170, 175, 176, 178 Laurentia 6 brachiopods 63 continental shelf 11 faunas 14, 49 Grenville belt 88 location 213 microplate detachment 75 rifted margin 18, 19, 41, 73 terranes 1, 5, 78 wander path 57 Laurentia-Antarctica connection 219 Laurentia-Gondwana collision 70, 76, 77, 219, 224, 261,277 interactions, 75 rifting 74 suture 177, 178 lavas, rhyolitic 73 Leptella (Leptella) 64 LepteIla (Petroria) 64, 69 Leptellinae 63 leucogranites 155, 197, 198, 207, 208, 290

leucosomes El PilSn 193 Sierra de Fiambalfi 301,302 Sierra de San Luis 242 Sierras de CSrdoba 184 L~vis Shale 63 lherzolite 183 LIL elements 198 Limay Mahuida 85 limestones, deep-water 38, 41 lineations, Sierra de Fiambalfi 307, 310 lit-par-lit injection 302, 321 lithospheric loading 76 lithostratigraphy, pre-Carboniferous 37 Llano 88, 177 Llanoria concept 86 87 Llanvirn brachiopods, Precordillera 61 Loma Ancha granite 185 Loma Colorada 223 Los Alanices 335 Los Azules Formation 94, 109, 120 Los Blanquitos block 22 Los Llanos-Ulapes batholith 290, 292, 344, 350-352 crustal component 365 major elements 354 QAP diagram 351 rare earth element spectra 356 Rb v. Yb + Ta plot 356 trace elements 355 Los Ramblones 223 Los Sombreros Formation continental slope 91 debris flows 21 fault scarp 48 olistoliths 45, 89, 90, 99 Los Tflneles clasts 46 metasediments 193 olistolith 41, 88, 89, 90 phyllites 210 Rb-Sr data 204, 358 shear zone 185 Lower Turbidite System 128, 129 geochemistry 134-135 lutites 285 mafics analyses 162 arc environments 163 composition 167 crystallization age 19 MORB-type 24 Pie de Palo complex 160, 172 magmas cogenetic 362 water content 115 magmatic arcs 3, 136, 227 magmatic development 113-114, 117 magmatic differentiation 334 magmatic evolution Sierra de San Luis 336 Sierras de CSrdoba 209-210 magmatic flow 348

375

376 magmatic suites, isotope signatures 182 magmatism Miocene 1 Sierras Pampeanas 52 major elements, Los Llanos-Ulapes batholith 354 Malanzfin Formation 262 Malvinas/Falkland Plateau 71,220 Mamil Choique 225 mantle, enriched 223 mantle doming 293 mantle upwelling 292 maps, regional 2 Marathon, bentonites 21 Marathon Formation 93 Marathon/Solitario region 85-86, 87, 92 asymmetry 100 basin evolution 97, 100 Maravillas Formation 94 marble, Pie de Palo 175, 178, 228 margin deformation 14 marginal plateaux 51 marine depositional environments 39 mass transfer 223 mass-flow deposits 45, 47, 51 mass-wasting deposits 25 Maurice Ewing Bank 14 Mediterranean faunas 71 mega-breccias 45, 51, 91 Mejillones complex 211 Mendoza carbonate rocks 85 thrust-fold belt 12 M6rida Andes 76 Merlinia 66 Mermela Group 261 Mesdn Group 210, 327 meta-andesites 88 meta-bentonites 85, 90, 99 see also bentonites meta-greywackes 285 meta-rhyolites 88 meta-texites 184 metabasites, E1 Diquecito 192 metagabbro 167, 222 Ouachita belt 25 metamorphic complex, Sierra de Los Llanos 348 metamorphic grade Pie de Palo complex 146, 159, 160 Sierra de Fiambalfi 307 Taconian 227 metamorphism carbonate platform 50 and deformation 350 Sierras de C6rdoba 185 temperatures 306, 308, 312 metasediments Los Tfineles 193 Sierras Pampeanas 210, 353, 356 metasomatism 112 metavolcanic rocks 236 mica-schists, Sierra de San Luis 327 Micaschist Group 236, 241,253 microcontinent models 15, 25 brachiopod evidence 59, 68, 72

INDEX microptates 26, 59 microprobe analyses 191, 194, 204 mid-Iapetus ridge 18 migmatisation 190, 228 Sierra de Fiambal/t 301 304 Sierra de San Luis 236, 240 migmatites 160, 183, 193 Millbrig beds 120, 121 Mina Gonzalito, age 225 Mina Gonzalito Yaminu6 complex 225 mineralogy, K-bentonites 111 Miocene lavas, xenoliths 5 Mississippi Valley trough 18, 19 mobile elements 160 Mogote Corralitos 161, 163, 172, 173 molasse sediments 52 Molles Formation 66 monazite 202, 203, 205, 266 monazites dating 271-272, 312-313 Tera-Wasserburg plots 271 Monorthis 64, 66 Montoya Formation 95 Monument Spring Member 93 monzogranites 183, 195, 198, 328, 335, 364 MORB signatures, amphibolites 212 muscovite 130 Ar/Ar dating 152-153, 154 mylonites Devonian 253 Eastern Basement Complex 21,247, 248, 251 La Escaleritla granite 332 Quebrada Gauyaupa-Lima 174 Riocito 249 Sierra de Chepes 348, 349 Sierra de Fiambalfi 313 Sierra de Paganzo 289 Sierra de Pie de Palo 154, 155, 160, 169 Sierras de C6rdoba 185 Western Basement Complex 245 myrmekitic texture 331,333, 350 Mytton Flags 66 N-MORB signatures 171, 173, 183, 214 Nanorthis 66 Narvfiez pluton 290 nautiloids 42, 90 Nazca plate, flat-slab zone 1 Negro Peinado Formation 285, 293 Nemagraptus gracilis zone 99 Neseuretus Fauna 66 Newfoundland, trondhjemites 228 NEWMOD program 112 Nikizanga block 163, 172 Niquivilia Zone 63, 64 Nogoli Metamorphic Complex 236 norite 222 North America, carbonate rim 20 North America-South America interaction 219 North Atlantic Region 76 North Patagonian Massif 220, 224, 225, 228 Nothorthis 63, 69 Notorthisina 66

INDEX Notus Ocean 77 Nuia 98 l~ufiorco pluton 290, 362 O. evae Zone 64 Oanduporella 67 obduction, ocean crust 100, 214 Occidentalia terrane 5, 53, 70, 219, 220, 229 ocean closure 72, 212, 222 oceanic crust, in clastic wedge 15 oceanic currents 59, 70 Ocloyic orogeny 11, 14, 21, 25, 27 contractional tectonics 100 Precordillera collision 123 Sierra de Fiambalh 319 Sierras Pampeanas 327 Ocoee Group 46 Oe. evae Zone 108, 122 Oepikoides 65 Ogof Hen Formation 64 Oklahoma aulacogen 92 Oklahoma-Alabama transform 85 Oldhamia 74, 222 Olenid biofacies 72 Oligorthis 64, 69 olistoliths 21, 25, 38, 41 Dagger Flat Sandstone 92, 93 Los Sombreros Formation 45, 89, 90, 99 Los Tuneles 41, 88, 89, 90 Woods Hollow Formation 94 olistostromes 91, 93 Olta 275, 349 oncolites 39, 43 onlap 41, 48 Onychopyge 66 oolites 39, 42 ophiolites Famatinean 19 Pie de Palo 146, 178 Sierras de C6rdoba 212 western Precordillera 25 Ordovician outcrop map 12 Precordillera docking 11, 12 Oriskany Formation 77 orogenic cycles 1 3 orogenic events 4 orogenic material, recycled 130 Orthidiella 64 Orthidium 63 orthogneiss felsic 261 La Puntilla 305, 306, 31 l, 313 metaluminous 195, 209 peraluminous 183 Sierra de Umango 220 Ouachita embayment 18, 88 detachment from 11, 15, 75, 177, 219 oceanic crust 100 passive margins 14, 19 reconstruction 99 structure and sedimentation 85 Ouachita fold belt, pyroclastic flows 85

Ouachita rift 18 Ouachita trough 228 outer shelf deposits 93, 96 overstep, and rift drift completion 19 P. elegans (P. suecicus) Zone 122 P. serra Zone 65 P - T conditions Rio de Susquia 193 Sierra de Fiambal'~ 307, 311 312 Sierra de Paganzo 364 Sierras de C6rdoba 193, 195, 210, 212 Pacatala 352 packstones 38, 39, 42 Paganzo basin 35 Paganzo Group 327 paired rift margin hypothesis 68 palaeo-biogeography 7, 44 palaeocurrents 50, 128, 211 clastic rocks 285 palaeogeographical hypotheses 58, 67 palaeogeographical models 67 palaeogeography, early Palaeozoic 89 palaeolatitude, Carolina terrane 76 palaeomagnetism Famatina 73 Precordillera 14 palaeosomes 190 Paliella 285 Pampa de Las Invernadas 245 Pampean cycle 2, 3, 222, 235, 327 Pampean Orogeny 181,210, 211,259 collisonal event 7 extent 214 metamorphism 210, 251,275-277 pluton emplacement 318, 321 Pampean terrane 57, 212, 251 Pampia terrane 21, 73 PAMPRE Project 181 Pan-African cycle 285 Pancanta 241,327 geological map 328 Pangaea break-up 68 Parabolina ( Neoparabolina) frequens 66 Paradoxides 68, 74, 76 paragneiss, Sierra de Fiambalh 301 Paran~-Etendeka flood basalts 3 ParomaIomena polonica 65 Paso del Rey 242, 250, 332, 333, 335 passive margins Appalachian 16 carbonate successions 15, 226 Cuyania 97 destruction of 214 extension 24, 52 major dement evidence 138 Ouachita embayment 19 Pampean cycle 3, 211,213 recycled sources 137 rifting 27, 51 trace element evidence 139 140 Patagonia plate 73 Patagonian massifs 5

377

378 Pb-isotopic signature Grenvillian 88, 145 Laurentian crust 92 Pb/Pb ages La Puntilla orthogneiss 305 Pie de Palo complex 146, 163, 175 Peacock index 353 Peggy Ridge 175 pegmatites 185, 236, 241,242, 250, 334 peneplain, pre-Andean 147 Penobscot orogeny 74, 228 pentamerids 63, 66 Perc6 76 petrography, turbidites 129-131 Phanerozoic cycles, table 3 phenocrysts 111, 112, 113 Phyllite Group 236, 241 structures 243-245 phyllites Los Tfineles 210 Sierra de San Luis 327 phyllonites 185 Pichanas complex 275 Pie de Palo complex 146, 160, 177 suturing 178 Piedmont zone 226, 227, 229, 231 piemontite 286 pillow-lavas Alcaparrosa Formation 24, 46, 95 rift-related 25, 49, 52, 99 planktonic stages, brachiopods 59 Planolites 285 plants, lycophyte 77 plate configuration 19-20, 100 plate interactions 52 plate rotation 100 platform area 37, 47 platform successions 72 Platillaenus 73 Platystrophia 64 Plectothyrella crassicosta 65 Plethopeltis obtusus 63 Plicatolina 66 Pliomeridius 66 plutonic belts 3, 182 plutonic rocks, Famatina System 288-290 plutons calc-alkaline 225 fracture-controlled 223 Pocho Phyllites 193 polymetamorphism 182 Pomatotrema 64 Pon6n Trehue basement 146 carbonate platform 22, 98, 143 Formation 91 Lower Ordovician 90 in Precordillera 21 Portezuelo del Tontal Formation 46, 91, 92 post-collisional cycle 14 Potomac orogeny 226 Potrerillo Block 22, 223 Prague basin 66 Prantlina 66

INDEX pre-Famatinian cycle, Sierra de San Luis 238-240 Precambrian-Cambrian boundary 212 Precordillera accretion 75 Argentina 1 collision 123 docking 11, 12, 78, 143, 277 drowning 20 East-West differences 16 evolutionary models 25, 27 as exotic terrane 3, 5, 6, 11 hypotheses 14 location map 13 origins 219 separation 18, 20, 27, 78 stratigraphic evidence 15 structural units 36 wander path 14, 25 width 16 prehnite 286 Pringles Metamorphic Complex 236, 277 Productorthis 64 proto-Andean margin, reconstruction 213 Protopeltura 66 provenance, discriminant plot 138 proximal facies 285 Pulchrilamina 98, 100 pumpellyite 286 Puna clastic rocks 129, 223 collision 25 connection with Famatina System 73, 128 magmatic arc 75, 127, 219 magmatic rocks 3 Plateau 297, 318 stratigraphy 128 thrusting 224 Turbidite Complex 128 Puncoviscan ocean 212, 214 Puncoviscana Formation 74, 75, 210, 211 age 211 basement 318 clastic sediments 213 equivalence 236, 251,275, 297 granitoids in 227 ichnofossils 222 Puncoviscana Trough 259 Punta Negra 146, 155 Punta Negra Formation 50 pyroclastic flows Fort Pefia Formation 93, 97 Ouachita fold belt 85 pyroxene 130 QAP diagram Famatina plutons 288 Los Llanos-Ulapes batholith 351 Sierra de San Luis 329 quartz gabbro 328 quartz grains analyses 116, 117 optical measurements 114 Quebrada de Los Saltitos 108 Quebrada de Potrerillos 109

INDEX Quebrada de Talacasto 109 Quebrada Grande 161 Quebrada Guayaupa-Lima 173 Quebrada Las Piriquitas 175, 178 Quebrada Las Plantas 109 Quebrada Nazareno 109 Quebrado del Gato 163, 174, 176 Quebrado del Molle 172 Quihnes 297 Quilpo Formation 275 Quines 236, 243, 250 R1-R2 diagrams 331 radiometric dating, see also Ar/Ar, K/Ar, Pb-Pb, U/Pb rafting, brachiopod larvae 59 Raleigh-Kiokee belt 226 ramp facies 23, 41 rare earth element spectra bentonites 117 Los Llanos-Ulapes batholith 356 Pie de Palo complex 167, 168, 169, 170, 174 Puna turbidites 136, 137 Ratones Granite 300, 313-317 R b - K diagram 137 Rb-Sr dating Achaia batholith 184 Bemberg tonalite 241,329 Cerro Toro granite 289 Chepes 359 Colohuincul 224 La Escaterilla granite 332 Sierra de Pie de Palo 161, 163, 220 Sierra de San Luis 333, 335 Sierras de C6rdoba 200-202, 204 Sierras Pampeanas 261,357-358, 358, 363, 364 Realito 241 receptaculids 42, 90 red-beds 17, 88 Redonia 66 reef associations, San Juan Formation 42 43, 90 regional metamorphism 313 relaxation, post-collisionaI 24 Renca 335, 336, 338 retro-arc environments 127 retrograde mineral growth 315, 317, 321 reverse faulting 250 Andean 1 Laramide 297 Neogene 318 reworking, heavy minerals 131 rhyodacite 131,286 rhyolite dykes 329 Famatina System 286 Gondwanian cycle 3 Laurentian margin 92, 93, 97 Puna arc 128 rift facies, late Precambrian 74 rift valleys 49 rift-drift succession 46 rift-drift transition 19, 27, 51, 213, 226 rift-onset unconformities 49, 52

rifting asymmetric 19 back-arc 175, 176 diachronous 15 evidence for 17 20, 27 initiation 68, 97 intracontinental 41, 51 Rinconada Formation 49, 50 Rio Cachiri Group 77 Rio Chico complex 225, 228 Rio Claro 241,245 Rio de La Cafiada Honda 250 Rio de la Carpa 242, 332, 333, 334 Rio de la Plata 211,212, 251,259, 277, 285 Rio de Suquia 185, 187, 197, 202 Rio Doce Orogeny 3 Rio Guzman Shear Zone 251,264, 278 Rio Quinto 245, 330 Rio Salado 225 Rio San Juan 49 Rio Sassito 47, 49 Rio Turbio, shear zone 247 Riocito, mylonites 249 Rochambeau Bank 175 Rodeo Viejo 335 Rodriguez Tank Sandstone 93 Rome (Waynesboro) Formation 18 Rosslare terrane 72 Rugostrophia 64 Sacabamaspisjanvieri 65, 70 Salopia 67 sample descriptions, Sierras Pampeanas 262-266 San Agustin 198, 208, 213, 290 San Carlos massif 195 San Isidro block 22 San Juan 88 San Juan Formation 42-44, 73, 90 reef association 98 volcanism 93 San Juan Limestone bentonites in 108, 109, 111 brachiopods 57, 60, 64 top of 120 San Luis Formation 261,277 San Rafael 36, 98 Grenvillean 44, 85, 222 San Rafael block 12 Santa Cruz de la Sierra 352 Santa Marta Massif 76 Santa Rosa 195, 208, 236, 243, 250 Santa Victoria Group 66 Sassito Formation, section 22, 23 Sauk unconformity 94, 95, 121 Sc-Th-Zr/10 plot 139 Scandinavian Caledonides 74 schists, Pie de Palo complex 173-174 Schizophoria 67 Scoto-Appalachian Province 65 Scottish-Irish Caledonides 74 Sea of Japan 98 sea-floor spreading, onset of 19 sea-level curve 45

379

380

INDEX

sea-level fluctuations 20, 22 global 43 seamounts 172, 175 section, block-faulting 24 sediment thickness 48, 49 sedimentary basins 64 sedimentary sequences, Precordillera 88 sedimentary units, overlapping relationships 12 sediments, Cambro-Ordovician 36 segmentation Andean 12 fault-controlled 92, 97 Sepulturas Formation 66 sequence boundaries type-1 40, 41, 42 type-2 40, 42 sequence stratigraphy 35 La Silla Formation 52 San Juan Formation 43 Zonda Formation 40 sericite 249, 250 serpentinite Ouachita thrust belt 25, 100 Pie de Palo complex 160, 163 Sierras de C6rdoba 183 shales black 228, 285 graptolitic 24, 72 phosphatic 93 shallow shelf depositional environments 39 shear zones Rio Turbio 247 Sierras de C6rdoba 184-185 Sierras Pampeanas 261,278 ultramafics in 227 Shimanto belt 212 Shin Brook Formation 66, 73, 74, 76 shortening, collisional 21 shoshonitic lavas 175, 178, 290 SHRIMP U-Pb analyses 6, 181,202, 204, 205, 266, 329, 361 Shumardia 66 Sierra Chica 185 Sierra Chica de Zonda 38, 49, 88, 177 Sierra de Ancasti 222, 229, 253, 261,318 Sierra de Chepes 251,275, 278, 336, 343 anatexis 350 andalusite zone 349 biotite zone 349 cordierite zone 349 geological map 263 granites 344 granodiorite 348 section 348 Tuani Granite 352 Sierra de Copacabana 285 Sierra de E1 Gigante 253 Sierra de Famatina 3, 60 biogeography 65-66 Sierra de Fiambalfi 7, 222, 229, 290, 297 amphibole barometry 303 block movement 314-315, 319 Cambrian 301-306 Carboniferous 315-317

correlation 319 geological map 301 geological outline 300 metamorphism 314 Mid-Palaeozoic 313-315 migmatisation 301 Ordovician 306-311 P - T conditions 307, 311-312 Precambrian 300 301 regional geology 318-319 sample location map 309, 316 structure and metamorphism 302, 306, 317 structure orientation 305, 310, 311 tectono-thermal events 317 textures and structures 304 thermo-barometry 311 Sierra de Guayaguas 253 Sierra de La Huerta 253, 336 Sierra de la Invernada 108 Sierra de la Macarena 68, 76 Sierra de Los Llanos 251,343 analyses 346-347 metamorphic complex 348 Sierra de Malanzfin 251,343 Sierra de Maz 229 Sierra de Paganzo 285, 289, 365 P - T conditions 364 Sierra de Paimfin 223, 285, 290 Sierra de Perijfi 76, 77 Sierra de Pie de Palo 6, 7, 85, 143 basement rocks 87, 98, 143, 159, 161 ductile deformations 147 geological map 145, 161 Grenvillian 220 map 160 marble 175, 178, 228 ophiolites 146 Rb-Sr dating 161, 163, 220 Sierra de Quilmes 318 Sierra de San Luis 7, 223, 235 block diagram 238, 246 crustal thickening 337 early intrusions 241 evolution 255 Famatinian cycle 251-255 geological map 237, 265, 326 granitoids 325 K-At ages 251 magmatic evolution 336 metamorphism 327 post-Famatinian magmatism 243-251 pre-Famatinian cycle 238-240, 251 profile 244, 254 QAP diagram 329 regional metamorphism 250~51 San Luis Formation 261 Sierra de Talacasto 95, 96 Sierra de Ulapes 343 Sierra de Umango 5, 220 Sierra de Valle F6rtil 156, 229, 253, 336 Sierra de Villicum 49, 88, 177, 228 Sierra Grande de C6rdoba 348 Sierra Norte de Cdrdoba 262, 277

INDEX Sierras de C6rdoba 181,251 age histogram 210 back-arc environments 214 concordia plots 206 detrital pattern 275 event sequence 183 geochronology 209 geological map 182, 186, 264 Harker diagrams 196 isochron plots 207 K A r data 187 lithology 182-184 magmatic evolution 209 210 ophiolites 212 P-Tconditions 193, 195 Tasa Cuna Formation 261 tectonothermal evolution 209-210 volcanic arc 222 Sierras Pampeanas 1, 3, 7, 12 analytical methods 266-267 Ar-Ar dating 147-149, 222 basement 36, 143, 159, 253, 261 Central Batholithic Belt 184 depositional ages 275 evolution 220, 228 geochronology 356-361 geological map 260, 344, 345 granitoids 343 intrusives 262 magmatism 21, 52, 212 metamorphism 261 metasediments 210 orogenic events 276 sample descriptions 262-266 shearing 261 southern 259 tectonic province 297, 325 tectono-thermal events 318 Th-Pb dating 267 U-Pb dates 267 ultramafics 261 zircon ages 270 sillimanite fibrolitic 241,242, 243, 333 Sierra de Fiambal~ 305, 306 siltstone association 39, 88 Skenidioides 66 slope apron 91 slope sediments 41 slumped beds 41 Sm Nd analyses Colohuincul 224 Sierra de Fiambal~ 312 Sierra de Los Llanos 359 Sierras de C6rdoba 203, 208 Sierras Pampeanas 363, 364 Soldano Member 41 South American Plate 219 South Atlantic, opening 3 Southern Iapetus 57 reconstruction 69 Southern Proto-Atlantic 57 space photographs 87, 89 sponge spicules 41, 42, 89, 90

spreading ridges hydrothermal activity 172 Quebradas del Quemado 169 remnants of 14 and submarine volcanism 24 Sr/Sr ratios 207, 361 starved-basin sediments 72 staurolite 240, 250 storm reworking 42 stratigraphic column 91 stratigraphic section 17 stratigraphy post-collision 21 Puna basin 128 stromatolites 40, 90, 175, 178 stromatoporoids 42, 90, 98 subduction 52, 59, 73, 75, 122 Early Cambrian 213 Famatinian 213, 214, 251,277, 292, 344 metamorphism 227 subduction zone 160, 177 subsidence controls on 17, 19, 41 plate configuration 20 Puna basin 128 thermal 52, 72 subsidence analysis 19, 20 supersequences 37 38 A 38-41 B 41-44 C 44-49 D 49-50 supra-subduction !75 supracrustal rocks, foreland 5 Suri Formation 66, 73 syn-collisional arc 228 syn-collisional cycle 14 syn-collisional suture 229 syn-rift successions 18 Syntrophia 63, 69 systems tracts highstands 39, 43 shelf-margin 43 transgressive 39, 43 Taconian orogeny Appalachians 225-227 continental collision 27, 219, 220, 228 deformation 156 metamorphic regime 223 origin of 11, 25 terrane accretion 74, 76, 122 Taconian-Famatinian belts, fit 229, 230 Taffia 64 Talacasto 73, 108, 111 Talacasto Formation 50 Tama Gabbro 352, 353, 362, 363, 365 Tamboreo 241,327, 328, 329 Tamengo Formation 211 Tandilia 111 Tasa Cuna Formation 261 Tastil 228 tectofacies sub-divisions 14

381

382

INDEX

tectonic discrimination major elements 137-138 trace elements 138-139 tectonic discrimination diagram 117 tectonic events 44 tectonic evolution, Cambrian 211-212 tectonic fields, granitoids 331, 333 tectonic overprinting 127 tectonic reworking 16 tectonic settings, major element evidence 138 tectonic tracer model 11, 25 tectono-sedimentary environments 47 tectono-sedimentary evolution 51 tectonothermal evolution, Sierras de C6rdoba 209-210 tempestites 39, 41, 43 tephra 112 Tepuel basin 71 Tera-Wasserburg plots 271, 273 terrane accretion 76 terrane hypothesis 11 terrane limits 13 terrane map 144 terranes, Laurentian 1 Tethyan Realm 71 Texas Plateau 59, 70, 71,235 Texas promontory 19, 71 Texas transform 85 Th-Pb dating analytical methods 266 monazites 271-272 Sierras Pampeanas 267 Th/U v. age plots 272 thermal doming 19 thermal gradient 350 thermo-barometry 182, 308 third-order cycles 38, 39, 40, 42 thrombolite mounds 40, 42, 90 thrust-fold belt, Precordillera 12 thrusting eastern Precordillera 21 Potomac orogeny 226 Puna 224 Sierra de Pie de Palo 147, 155, 177 Sierras de C6rdoba 185 Tilacaric orogeny 75 TiO2 v. Fe/Mg diagram 192 Tioga Bentonite 77 Tissintia 65, 67 Titicaca Group 71 Tofua arc 175, 176 tonalites t54, 195, 222, 241 chemistry 329-330 Famatinian cycle 251,327 Las Verbenas 243 Toquima-Table Head Realm brachiopods 57, 62, 64, 65, 71 passive margin 72 Tourmakeady Limestone 63, 64, 66, 73 tourmaline 130, 197 trace elements immobile 131 Los Llanos-Ulapes batholith 355 tectonic discrimination 138-139 trace fossils 211

trachyandesite 93, 286 trachyte 92, 93 transfer faults 15 transform faults 18, 85, 100, 219 transgression, post-rifting 17, 19 transpressive collision 229, 245 transtensional separation 98, 227 Trapiche Formation 21, 25, 47, 60 Treiorwerth Formation 64 trilobites agnostid 41, 89 cladistic analysis 65 correlation 57, 62 cyclopygid 66 deep-water 43 Laurentian 14, 35 Olenellid 19, 38 Tremadoc 66, 285 trondhjemite-tonalites 183, 185, 198, 207, 208 trondhjemites, Newfoundland 228 Tuani Granite 352, 353 Tucavaca Group 211 tufts, lapilli 128 turbidites 45, 50, 89, 91 framework minerals 128 geochemistry 131 Pampean 210, 285 petrography 129-131 provenance 127, 131 137 sample descriptions 128 129 weathering 131, 137 TWEEQU program 185 U. austrodentatus Zone 108, 109 U-Pb dates analytical methods 266, 298 Arequipa-Antofalla craton 220 Bemberg tonalite 329 bentonites 85, 109 Chepes Granodiorite 360, 361, 362 Eastern Basement Complex 240 granites 6 Grenvillean 87, 161 igneous zircons 273 La Escalerilla granite 243 monazites 271-272 Narvfiez 290 Pie de Palo complex 146, 163 Sierra de Fiambal/t 299, 300, 302, 312, 316 Sierra de San Luis 241 Sierras de C6rdoba 184, 198, 199, 202 Sierras Pampeanas 267, 359 xenoliths 5, 144 ultrabasics, Sierras de C6rdoba 183 ultramafics 160, 162, 163 in shear zones 227 unconformity, Sauk 94, 95, 121 underplating 213 underthrusting 212 unrelated margin hypothesis 67 68 Upper Turbidite System 128, 129 geochemistry 136

INDEX Venezuelan basin 77 Virgiliana orogeny 229 Virorco-Las Aguilas 229, 236, 247 Volcancito Formation 66, 285, 293 volcanic aprons 96, 128 volcanic arcs 20, 68, 72 Famatina System 286 Puna basin 127, 131,223 volcanic glass Fort Pefia 96 phenocrysts 113 turbidites 130 volcanic islands 68, 72 volcanic vents locations 117 marine 94 volcaniclasts 92, 93, 122, 128, 285 volcanism acid 293 arc type 20 explosive 117 Famatina System 286-288 Marathon 93 Precordillera 111 Volcanosedimentary Succession geochemistry 132--133 Puna basin 128, 129 wackestones 41, 42, 43 wander path Laurentia 57, 76 Precordillera 14 water content, magmas 115, 334 water temperature, and brachiopod dispersion 59 weathering, turbidites 131, 137, 140 websterite 183

wedges, volcaniclastic 72 Western Basement Complex 241,243 western basin 37 Western Iapetus, closure 71 Western Sierras Pampeanas assimilation 293 Espinal Formation 283 granitoids 122 Grenvillean 3, 5, 12 limit of Occidentalia 220 Sierra de E1 Gigante block 253 subduction and magmatism 100 whole-rock analyses, K-bentonites 118-119 Wilson Cycle 70 Woods Hollow Shale 93, 94, 97, 99 X-ray diffraction patterns 112 xenoliths Grenvillean 36, 144, 220 Miocene lavas 5 Y-Sr/Y diagram 197 zircon, Puna turbidites 130, 131 zircon ages comparison with monazite ages 272-273 Devonian 50 Famatinian system 95 Grenvillean 5 meta-bentonites 85 Sierras Pampeanas 270 zircons, cathodoquminescence images 269 Zonda Formation 40, 90 Zondarella communis 98, 100 Zr/Sc v. Th/Sc plot 137

383

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Geoarchaeology: Exploration, Environments, Resources (Geological Society Special Publication, No. 165) (Geological Society Special Publication, No. 165)

The Proto-Andean Margin of Gondwana (Geological Society Special Publication No. 142)

Geoarchaeology: Exploration, Environments, Resources (Geological Society Special Publication, No. 165) (Geological Society Special Publication, No. 165)

Geological Data Management (Geological Society Special Publication)

Terrane Processes at the Margins of Gondwana (Geological Society Special Publication No. 246)

Deformation of Glacial Materials (Geological Society Special Publication No. 176)

West Gondwana: Pre-Cenozoic Correlations across the South Atlantic Region (Geological Society Special Publication No. 294)

The Making of the Geological Society of London (Geological Society Special Publication No. 317)

Australian Landscapes (Geological Society Special Publication 346)

Martian Geomorphology (Geological Society Special Publication 356)

Communicating Environmental Geoscience - Special Publication no 305 (Geological Society Special Publication) (No. 305)

Sustainable Groundwater Development (Geological Society Special Publication,)

Early Precambrian Processes (Geological Society Special Publication)

Phanerozoic Ironstones (Geological Society Special Publication)

Slope Tectonics (Geological Society Special Publication 351)

Volcanic Degassing (Geological Society Special Publication No. 213)

Fractured Reservoirs (Geological Society Special Publication no 270)

Confined Turbidite Systems (Geological Society Special Publication No. 222)

Geomaterials in Cultural Heritage (Geological Society Special Publication No. 257)

Hydrothermal Vents and Processes (Geological Society Special Publication No. 87)

Climates: Past and Present (Geological Society Special Publication No. 181)

Hydrocarbons in Crystalline Rocks (Geological Society Special Publication No. 214)

Gas Hydrates: Relevance to World Margin Stability and Climatic Change (Geological Society Special Publication No.137)

Fractal Analysis for Natural Hazards - Special Publication No 261 (Geological Society Special Publication)

Palaeozoic Reefs and Bioaccumulations: Climatic and Evolutionary Controls - Special Publication no 275 (Geological Society Special Publication)

Devonian Events and Correlations : Special Publication no 278 (Geological Society Special Publication)

Geological Storage of Carbon Dioxide (Geological Society of London Special Publication No. 233)

The Physics of Explosive Volcanic Eruptions (Geological Society Special Publication)

Non-Marine Permian Biostratigraphy and Biochronology - Special Publication No. 265 (Geological Society Special Publication)

Structurally Complex Reservoirs - Special Publication no 292 (Geological Society Special Publication)

Building Stone Decay: From Diagnosis to Conservation - Special Publication no 271 (Geological Society Special Publication)

The Evolving Continents: Understanding Processes of Continental Growth - Special Publication 338 (Geological Society Special Publication)

The Proto-Andean Margin of Gondwana (Geological Society Special Publication No. 142)

Geoarchaeology: Exploration, Environments, Resources (Geological Society Special Publication, No. 165) (Geological Society Special Publication, No. 165)

Geological Data Management (Geological Society Special Publication)

Terrane Processes at the Margins of Gondwana (Geological Society Special Publication No. 246)

Deformation of Glacial Materials (Geological Society Special Publication No. 176)

West Gondwana: Pre-Cenozoic Correlations across the South Atlantic Region (Geological Society Special Publication No. 294)

The Making of the Geological Society of London (Geological Society Special Publication No. 317)

Australian Landscapes (Geological Society Special Publication 346)

Martian Geomorphology (Geological Society Special Publication 356)

Communicating Environmental Geoscience - Special Publication no 305 (Geological Society Special Publication) (No. 305)

Sustainable Groundwater Development (Geological Society Special Publication,)

Early Precambrian Processes (Geological Society Special Publication)

Phanerozoic Ironstones (Geological Society Special Publication)

Slope Tectonics (Geological Society Special Publication 351)

Volcanic Degassing (Geological Society Special Publication No. 213)

Fractured Reservoirs (Geological Society Special Publication no 270)

Confined Turbidite Systems (Geological Society Special Publication No. 222)

Geomaterials in Cultural Heritage (Geological Society Special Publication No. 257)

Hydrothermal Vents and Processes (Geological Society Special Publication No. 87)

Climates: Past and Present (Geological Society Special Publication No. 181)

Hydrocarbons in Crystalline Rocks (Geological Society Special Publication No. 214)

Gas Hydrates: Relevance to World Margin Stability and Climatic Change (Geological Society Special Publication No.137)

Fractal Analysis for Natural Hazards - Special Publication No 261 (Geological Society Special Publication)

Palaeozoic Reefs and Bioaccumulations: Climatic and Evolutionary Controls - Special Publication no 275 (Geological Society Special Publication)

Devonian Events and Correlations : Special Publication no 278 (Geological Society Special Publication)

Geological Storage of Carbon Dioxide (Geological Society of London Special Publication No. 233)

The Physics of Explosive Volcanic Eruptions (Geological Society Special Publication)

Non-Marine Permian Biostratigraphy and Biochronology - Special Publication No. 265 (Geological Society Special Publication)

Structurally Complex Reservoirs - Special Publication no 292 (Geological Society Special Publication)

Building Stone Decay: From Diagnosis to Conservation - Special Publication no 271 (Geological Society Special Publication)

The Evolving Continents: Understanding Processes of Continental Growth - Special Publication 338 (Geological Society Special Publication)

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