The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics (Geological Society Special Publication)

The Neuquen Basin, Argentina A Case Study in Sequence Stratigraphy and Basin Dynamics

Geological Society Special Publications Books Editorial Committee BOB PANKHURST (UK) (CHIEF EDITOR Society Books Editors J. GREGORY (UK) J. GRIFFITHS (UK) J. HOWE (UK) P. LEAT (UK) N. ROBINS (UK) J. TURNER (UK) Society Books Advisors M. BROWN (USA) R. GIERE (Germany) J. GLUYAS (UK) D. STEAD (Canada) R. STEPHENSON (Netherlands) S. TURNER (Australia)

Geological Society books refereeing procedures The Society makes every effort to ensure that the scientific and production quality of its books matches that of its journals. Since 1997, all book proposals have been refereed by specialist reviewers as well as by the Society's Books Editorial Committee. If the referees identify weaknesses in the proposal, these must be addressed before the proposal is accepted. Once the book is accepted, the Society has a team of Book Editors (listed above) who ensure that the volume editors follow strict guidelines on refereeing and quality control. We insist that individual papers can only be accepted after satisfactory review by two independent referees. The questions on the review forms are similar to those for Journal of the Geological Society. The referees' forms and comments must be available to the Society's Book Editors on request. Although many of the books result from meetings, the editors are expected to commission papers that were not presented at the meeting to ensure that the book provides a balanced coverage of the subject. Being accepted for presentation at the meeting does not guarantee inclusion in the book. More information about submitting a proposal and producing a Special Publication can be found on the Society's web site: www.geolsoc.org.uk.

It is recommended that reference to all or part of this book should be made in one of the following ways: VEIGA, G. D., SPALLETTI, L. A. HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252. MCILROY, D. & FLINT, S., HOWELL, J. A. & TIMMS, N. 2005. Sedimentology of the tide-dominated Jurassic Lajas Formation, Neuquen Basin, Argentina. In: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 83-107.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 252

The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics EDITED BY

G. D. VEIGA Universidad Nacionald de La Plata-CONICET, Argentina

L. A. SPALLETTI Universidad Nacionald de La Plata-CONICET, Argentina

J. A. HOWELL University of Bergen, Norway

and

E. SCHWARZ University of Ottawa, Canada

2005 Published by The Geological Society London

THE GEOLOGICAL SOCIETY The Geological Society of London (GSL) was founded in 1807. It is the oldest national geological society in the world and the largest in Europe. It was incorporated under Royal Charter in 1825 and is Registered Charity 210161. The Society is the UK national learned and professional society for geology with a worldwide Fellowship (FGS) of 9000. The Society has the power to confer Chartered status on suitably qualified Fellows, and about 2000 of the Fellowship carry the title (CGeol). Chartered Geologists may also obtain the equivalent European title, European Geologist (EurGeol). One fifth of the Society's fellowship resides outside the UK. To find out more about the Society, log on to www.geolsoc.org.uk. The Geological Society Publishing House (Bath, UK) produces the Society's international journals and books, and acts as European distributor for selected publications of the American Association of Petroleum Geologists (AAPG), the American Geological Institute (AGI), the Indonesian Petroleum Association (IPA), the Geological Society of America (GSA), the Society for Sedimentary Geology (SEPM) and the Geologists' Association (GA). Joint marketing agreements ensure that GSL Fellows may purchase these societies' publications at a discount. The Society's online bookshop (accessible from www.geolsoc.org.uk) offers secure book purchasing with your credit or debit card. To find out about joining the Society and benefiting from substantial discounts on publications of GSL and other societies worldwide, consult www.geolsoc.org.uk, or contact the Fellowship Department at: The Geological Society, Burlington House, Piccadilly, London W1J 0BG: Tel. +44 (0)20 7434 9944; Fax +44 (0)20 7439 8975; E-mail: enquiries @geolsoc.org.uk. For information about the Society's meetings, consult Events on www.geolsoc.org.uk. To find out more about the Society's Corporate Affiliates Scheme, write to [email protected].

Published by The Geological Society from: The Geological Society Publishing House Unit 7, Brassmill Enterprise Centre Brassmill Lane Bath BA1 3JN, UK Orders: Tel. +44 (0)1225 445046 Fax +44 (0)1225 442836 Online bookshop: www.geolsoc.org.uk/bookshop The publishers make no representation, express or implied, with regard to the accuracy of the information contained in this book and cannot accept any legal responsibility for any errors or omissions that may be made. © The Geological Society of London 2005. All rights reserved. No reproduction, copy or transmission of this publication may be made without written permission. No paragraph of this publication may be reproduced, copied or transmitted save with the provisions of the Copyright Licensing Agency, 90 Tottenham Court Road, London W1P 9HE. Users registered with the Copyright Clearance Center, 27 Congress Street, Salem, MA 01970, USA: the item-fee code for this publication is 0305-8719/05/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN 1-86239-190-4 Typeset by Techset Composition, Salisbury, UK Printed by MPG Books Ltd, Bodmin, UK.

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Contents

Preface HOWELL, J. A., SCHWARZ, E., SPALLETTI, L. A. & VEIGA, G. D.

vii 1

The Neuquen Basin: an overview Geodynamic and tectonic evolution RAMOS, V. A. & FOLGUERA, A. Tectonic evolution of the Andes of Neuquen: constraints derived from the magmatic arc and foreland deformation

15

ZAPATA, T. & FOLGUERA, A. Tectonic evolution of the Andean Fold and Thrust Belt of the southern Neuquen Basin, Argentina

37

Biostratigraphy AGUIRRE-URRETA, M. B., RAWSON, P. F., CONCHEYRO, G. A., BOWN, P. R. & OTTONE, E. G. Lower Cretaceous (Berriasian-Aptian) biostratigraphy of the Neuquen Basin

57

Sedimentary geology and sequence stratigraphy in continental to shallow-marine deposits MclLROY, D., FLINT, S., HOWELL, J. A. & TIMMS, N. Sedimentology of the tide-dominated Jurassic Lajas Formation, Neuquen Basin, Argentina SCHWARZ, E. & HOWELL, J. A. Sedimentary evolution and depositional architecture of a lowstand sequence set: the Lower Cretaceous Mulichinco Formation, Neuquen Basin, Argentina

83 109

VEIGA, G. D., HOWELL, J. A. & STROMBACK, A. Anatomy of a mixed marine-non-marine 139 lowstand wedge in a ramp setting. The record of a Barremian-Aptian complex relative sea-level fall in the central Neuquen Basin, Argentina STROMBACK, A., HOWELL, J. A. & VEIGA, G. D. The transgression of an erg - sedimentation and reworking/soft-sediment deformation of aeolian facies: the Cretaceous Troncoso Member, Neuquen Basin, Argentina

163

Sedimentary geology and cyclostratigraphy in offshore deposits DOYLE, P., POIRE, D. G., SPALLETTI, L. A., PIRRIE, D., BRENCHLEY, P. & MATHEOS, S. D. Relative oxygenation of the Tithonian-Valanginian Vaca Muerta-Chachao formations of the Mendoza Shelf, Neuquen Basin, Argentina

185

SCASSO, R. A., ALONSO, M. S., LANES, S., VILLAR, H. J. & LAFFITTE, G. Geochemistry and petrology of a Middle Tithonian limestone-marl rhythmite in the Neuquen Basin, Argentina: depositional and burial history

207

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CONTENTS

SAGASTI, G. Hemipelagic record of orbitally-induced dilution cycles in Lower Cretaceous sediments of the Neuquen Basin

231

TYSON, R. V., ESHERWOOD, P. & PATTISON, K. A. Organic facies variations in the Valanginian-mid-Hauterivian interval of the Agrio Formation (Chos Malal area, Neuquen, Argentina): local significance and global context

251

Palaeoecology and palaeobiology MORGANS-BELL, H. S. & MCILROY, D. Palaeoclimatic implications of Middle Jurassic (Bajocian) coniferous wood from the Neuquen Basin, west-central Argentina

267

GASPARINI, Z. & FERNANDEZ, M. Jurassic marine reptiles of the Neuquen Basin: records, faunas and their palaeobiogeographic significance

279

LAZO, D. G., CICHOWOLSKI, M., RODRIGUEZ, D. L. & AGUIRRE-URRETA, M. B. Lithofacies, palaeoecology and palaeoenvironments of the Agrio Formation, Lower Cretaceous of the Neuquen Basin, Argentina

295

CORIA, R. A. & SALGADO, L. Mid-Cretaceous turnover of saurischian dinosaur communities: evidence from the Neuquen Basin

317

Index

329

Preface

The aims of this special publication are to present the geological history of the spectacular Neuquen Basin. It is envisaged that this book will act as both an introduction to the basin and also as a focus for recent developments in the long history of its study. Furthermore, we hope that the book goes further than just presenting the latest studies on a specific area. We have aimed to present an integrated case study in sequence stratigraphy, palaeontology and basin analysis, lessons from which have implications for systems worldwide. The concept of this book was born in the field in Argentina. We felt that there was a need to provide high-quality case studies that integrate different aspects of basin evolution and sequence stratigraphy, and we felt the Neuquen Basin provided such a dataset. However, it was apparent that, in spite of the excellent outcrops, the fascinating geology and the number of groups working on many different aspects of the basin, it was not well known within the wider geological community. Despite the remoteness of much of the region, many studies have been undertaken. Some of these have been driven by the prolific hydrocarbons found in the basin. Others capitalize on the basin's unique palaeontological record, while the region has also been used for detailed sequence stratigraphic and facies-based studies that utilize the excellent outcrops as analogues for subsurface reservoirs both within the basin and internationally. Structural studies on the fold and thrust belt are central to understanding the evolution of the basin and the Andean margin of Gondwana. The structural history of the basin records the change from Late Triassic extension through Jurassic and Early Cretaceous thermal subsidence to middle Cretaceous foreland basin subsidence followed by Andean compression and uplift. Most recent stages of the evolution include the emplacement and extrusion of a variety of igneous suites. The basin-fill succession includes deepmarine turbidite, hemi-pelagic and pelagic

systems; shallow-marine clastic and carbonate systems; evaporites; a variety of different fluvial systems; and several phases of aeolian deposition. Facies contacts such as aeolian sandstones packages within ammonite-bearing offshore shales, are testament to a dramatic relative sealevel history, much of which is linked to the tectonic evolution of the Andes. The succession also contains one of the world's most important Mesozoic fossil records. The vertebrate record includes numerous finds of marine and terrestrial reptiles, many of which are unique. Less spectacular, but of comparable importance, is the most complete southern hemisphere Mesozoic invertebrate record. We hope that this special publication will provide an insight into this amazing area. In compiling this collection of papers we have aimed to illustrate how all aspects of the geology are interlinked and how all should be considered together. In addition, our hope is that this publication will serve both as a stepping stone to the region for further study and as a more general case study in integrating the many aspects of basin evolution. The Neuquen Basin is a unique area, with fantastic outcrops and many remaining problems to be solved. We believe that it will continue to be used as an important field laboratory and training ground for the geologists of the future. The editors are indebted to the authors who have contributed with timely and excellent contributions. The editors also acknowledge the efforts of the various reviewers who gave their time to ensure the scientific and linguistic integrity of the final product. They are N. Bardet, J. Battacharya, R. Bersezio, R. Blakey, P. de Boer, L. Buatois, P. Cobbold, K. Curry Rogers, P. Doyle, W. Etter, J. Franzese, F. Fursich, M. Gruszczynski, A. Guierrez-Pleimling, A. Hartley, J. Hechem, M. Lamanna, C.O. Limarino, D. Loope, C. Morley, J. Mutterlose, E. Nichols, G. Plint, A. Ruffell, R. Scasso, T. Sempere, A. Tankard, A. Tripaldi, H. Weissert, H. Welsink, P. Wignall, B. Willis and A. Zamuner. They also thank those reviewers

viii

PREFACE

who decided to remain anonymous. The impetus for this publication was born out of the long-term relationship between Universities of La Plata and Liverpool, and as such we would like to acknowledge the important role played by Professor S. Flint. Angharad Hills at the

Geological Society played an essential role with her invaluable advice. Gonzalo Veiga Luis Spalletti John Howell & Ernesto Schwarz

The Neuquen Basin: an overview JOHN A. HOWELL1, ERNESTO SCHWARZ2, LUIS A. SPALLETTI3 & GONZALO D. VEIGA3 1 Centre for Integrated Petroleum Research, University of Bergen, Allegt. 41, N-5007 Bergen, Norway (e-mail: [email protected]) Department of Earth Sciences, University of Ottawa, 140 Louis Pasteur Pvt, Ottawa, Canada KIN 6N5 3 Centro de Investigaciones Geologicas, Universidad Nacionald de La Plata-CONICET, Calle 1 No. 644, B11900TAC, La Plata, Argentina Abstract: The Neuquen Basin of Argentina and central Chile contains a near-continuous Late Triassic-Early Cenozoic succession deposited on the eastern side of the evolving Andean mountain chain. It is a polyphase basin characterized by three main stages of evolution: initial rift stage; subduction-related thermal sag; and foreland stage. The fill of the basin records the tectonic evolution of the central Andes with dramatic evidence for baselevel changes that occurred both within the basin and along its margins. The record of these changes within the mixed siliclastic-carbonate succession makes the basin an excellent field laboratory for sequence stratigraphy and basin evolution. The 4000 m-thick fill of the basin also contains one of the most complete Jurassic-Early Cretaceous marine fossil records, with spectacular finds of both marine and continental vertebrates. The basin is also the most important hydrocarbon-producing province in southern South America, with 280.4 x 106 m3 of oil produced and an estimated 161.9 x 106 m3 remaining. The principal components of the hydrocarbon system (source and reservoir) crop out at the surface close to the fields. The deposits of the basin also serve as excellent analogues to reservoir intervals worldwide.

This paper aims to provide a brief introduction to the Neuquen Basin. It should provide a stepping stone for further reading and also for further studies. This paper also serves as an introduction to this Special Publication, which details the most recent work within the basin. The proposed goals of the Special Publication are as follows. • • • •

To present the Neuquen Basin as an integrated case study in sequence stratigraphy and basin analysis. To document the latest developments in vertebrate and invertebrate palaeontology. To consider the basin in the context of the structural evolution of the central Andes. To document the latest studies on specific stratigraphic intervals in a way that allows the reader to build up a complete picture of the basin fill and the way in which the various depositional systems have evolved through time.

•

To present specific studies from the basin that highlights concepts and models in sequence stratigraphy that are exportable to other systems.

Introduction to the Neuquen Basin The Neuquen Basin is located on the eastern side of the Andes in Argentina and central Chile, between 32° and 40°S latitude (Figs 1 & 2). It covers an area of over 120 000 km2 (Yrigoyen 1991) and comprises a continuous record of up to 4000 m of stratigraphy. This Late Triassic Early Cenozoic succession includes continental and marine siliciclastics, carbonates and evaporites that accumulated under a variety of basin styles (Fig. 3). The basin has a broadly triangular shape (Fig. 1) and two main regions are commonly recognized: the Neuquen Andes to the west,

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 1-14. 0305-8719/05/$15.00 © The Geological Society of London 2005.

2

J. A.hHOWELLETAL.

Fig. 1. Sketch map of the Neuquen Basin showing the approximate location (boxes and stars) of the contributions included in this publication. 1, Ramos & Folguera; 2, Zapata & Folguera; 3, Aguirre-Urreta et al.; 4, Mcllroy et al.;5, Schwarz & Ho well; 6, Veiga et al.; 7, Stromback et al.; 8, Doyle et al.; 9, Scasso et al.; 10, Sagasti; 11, Tyson et al.; 12, Morgans-Bells & Mcllroy; 13, Gasparini & Fernandez; 14, Lazo et al.; 15, Coria & Salgado.

and the Neuquen Embayment to the east and SE. The majority of the Basin's hydrocarbon fields are located in the Neuquen Embayment where most of the Mesozoic sedimentary record is in the subsurface and the strata are relatively undeformed. This is in contrast to the Andean region where Late Cretaceous-Cenozoic deformation has resulted in the development of a series of N-S-oriented fold and thrust belts (Aconcagua, Marlargiie and Agrio fold and thrust belts, Fig. 2) that provide excellent outcrops of the Mesozoic successions. During present times and throughout much of its history the triangular Neuquen Basin has been limited on its NE and southern margins by

wide cratonic areas of the Sierra Pintada Massif and the North Patagonian Massif, respectively (Fig. 1). The western margin of the basin is the Andean magmatic arc on the active western margin of the Gondwanan-South American Plate. This geotectonic framework and the highly complex history of the basin are largely controlled by changes in the tectonics on the western margin of Gondwana. The evolution and development of the basin can be considered in three stages (Fig. 3). 1.

Late Triassic-Early Jurassic: prior to the onset of subduction on its western margin,

THE NEUQUEN BASIN: AN OVERVIEW

3

Fig. 2. Major morphotectonic features of the Neuquen Basin and Andean Cordillera (Landsat image courtesy of Dr A. Folguera). Selected Cenozoic volcanoes are indicated by dotted lines. Inset shows image location in the Neuquen Basin.

this part of Gondwana was characterized by large transcurrent fault systems. This led to extensional tectonics within the Neuquen Basin and the evolution of a series of narrow, isolated depocentres (Manceda & Figueroa 1995; Vergani et al 1995; Franzese & Spalletti 2001). 2. Early Jurassic-Early Cretaceous: development of a steeply dipping, active subduction zone and the associated evolution of a magmatic arc along the western margin of Gondwana led to back-arc subsidence within the Neuquen Basin. This post-rift stage of basin development locally accounts for more than 4000 m of the basin fill (Vergani et al 1995).

3. Late Cretaceous-Cenozoic: transition to a shallowly dipping subduction zone resulting in compression and flexural subsidence, associated with 45—57 km of crustal shortening (Introcaso et al 1992; Ramos 1999Z?) and uplift of the foreland thrust belt. The final phase of Andean tectonism produced the uplift of the tightly folded outcrops in the western part of the area (Fig. 2). These outcrops expose a complete Mesozoic succession that includes a very wide variety of depositional settings. The lateral extent and spatial distribution of the deposits facilitates stratigraphic correlation and the tracing of regional unconformities. These outcrops have been used to understand

4

J. A.hHOWELLETAL.

Fig. 3. Chronostratigraphy, tectonic history and biostratigraphy of the Neuquen Basin. Lithostratigraphy is mostly after Legarreta & Gulisano (1989) and Legarreta & Uliana (1991). Only nomenclature of the Neuquen sector of the basin is depicted. Tectonic history after Vergani et al. (1995) and Franzese et al. (2003). Biostratigraphic resolution after Riccardi et al. (1999) (Jurassic), Aguirre-Urreta & Rawson (1997), Aguirre-Urreta et al (1999) (Early Cretaceous) and Casadio et al. (2004) (Late Cretaceous).

THE NEUQUEN BASIN: AN OVERVIEW

hydrocarbon reservoir systems both in the adjacent subsurface systems (Valente 1999; Vergani et al. 2002) and also worldwide (Brandsaeter et al 2005). The palaeontology of the Neuquen Basin is central to its global significance. The basin contains one of the most complete records of Jurassic and Cretaceous marine invertebrates. The completeness of this record has allowed the construction of accurate biostratigraphic charts for western Gondwana (Aguirre-Urreta et al 1999; Riccardi et al 1999). These charts allow excellent correlation and dating within the basin, and comparative correlation to faunas and successions from other parts of the world, for example North America and Thethys. The Mesozoic continental and marine reptile record of the Neuquen Basin is one of the most complete, varied and well preserved in the entire world. New theories with global implications on taxonomy, palaeobiogeography, palaeoecology and taphonomy merged from the study of these herpetofaunas (Gasparini 1996; Gasparini & Fernandez 1997; Gasparini et al 1997, 1999; Wilson & Sereno 1998; Sereno 1999). The Neuquen Basin has been the subject of numerous studies since the beginning of the 20th century. Prior to the 1960s early work included regional studies on the stratigraphy, palaeontology, biostratigraphy and structural geology (e.g. Weaver 1931; Groeber 1946; Herrero Ducloux 1946; De Ferrariis 1947; Groeber et al 1953). From the 1960s to the 1990s a concerted hydrocarbon exploration effort by YPF (the Argentinian National Oil Company), coupled with numerous academic studies, led to significant advances in the understanding of the basin. During this period the different structural styles were defined (Ramos 1978; Feehan 1984; Ploszkiewicz et al 1984), the biostratigraphic charts for the Jurassic and the Cretaceous were refined and updated (Riccardi et al 1971; Leanza 1973, 1981; Leanza et al 1977; Riccardi 1983), and the early schemes for the regional sequence and seismic stratigraphy were developed (Gulisano et al 1984; Mitchum & Uliana 1985; Legarreta & Gulisano 1989; Legarreta & Uliana 1991, 1999; Legarreta et al 1993). Since the early 1990s studies within the basin (including those presented in this Special Publication) have utilized the regional frameworks to address specific issues such as high-resolution sequence stratigraphic problems, detailed palaeogeographic and sedimentological studies of specific intervals, improved biostratigraphic charts and geochemical studies.

5

Geodynamic evolution The Neuquen Basin originated in the Late Triassic as a result of continental intraplate extension. During this period a series of extensional troughs were filled with volcaniclastic and continental deposits. During the subsequent growth of the Andean magmatic arc the basin became a back-arc system with widespread marine sedimentation. Acceleration of plate convergence during the Late Cretaceous produced partial inversion and the development of a retro-arc flexural system. This was associated with a progressive change from marine to continental sedimentation. The evolution of the Neuquen Basin is intimately linked to the development of the Neuquen Andes and the geometry of the subducting slab (Ramos & Folguera this volume). Late Trias sic-Early Jurassic sy nrift phase The Late Triassic-Early Jurassic margin of Gondwana in the vicinity of the Neuquen Basin lacks evidence for slab subduction. The tectonic system was dominated by a strike-slip regime subparallel to the western continental margin (Franzese & Spalletti 2001). In the area of the Neuquen Basin extension related to the collapse of the Gondwana Orogen produced a series of long, narrow half-grabens (Fig. 4A) that were filled by a complex array of clastic and volcaniclastic deposits associated with extensive lava flows (Franzese et al 2006) (Lapa Formation, Fig. 3, and equivalent units). Clastic deposits include alluvial, fluvial, shallow-marine, deltaic and lacustrine deposits (Franzese & Spalletti 2001). Fault growth, interaction and a transition to more regional subsidence during Early Jurassic times resulted in a more widespread lacustrine and shallow-marine facies distribution. Early Jurassic-Early Cretaceous post-rift phase During the Early-Middle Jurassic the subduction regime along the western Gondwana margin was initiated (Franzese et al 2003) and by the Late Jurassic the Andean magmatic arc was almost fully developed. Back-arc subsidence led to an expansion of the marine realm and flooding of the basin (Fig. 4B), which was connected to the proto-Pacific through gaps in the arc (Spalletti et al 2000; Macdonald et al 2003). Initially sedimentation was strongly influenced by the topography inherited from the underlying synrift systems (Burgess et al 2000; Mcllroy et al. this volume). After this initial

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J. A. HOWELLETAL.

Fig. 4. Schematic evolution of the Neuquen Basin from the Late Triassic to the Cenozoic. (A) Late Triassic-Early Jurassic, characterized by pre-subduction rifting in a series of narrow grabens. (B) Jurassic-Early Cretaceous, onset of subduction on the western margin of Gondwana and the early development of the Andean chain. The basin is a large triangular embayment periodically separated from the proto-Pacific by uplift and relative sea-level fall. (C) Late Cretaceous Andean uplift, development of a foreland thrust belt and basin. Much of the basin fill is non-marine, although periodic transgression from the Atlantic results in some marine intervals. Based on Vergani et al (1995), Ramos (1999&), Franzese & Spalletti (2001), Folguera & Ramos (2002) and Franzese et al. (2003). Original drafts courtesy of Dr J. Franzese.

THE NEUQUEN BASIN: AN OVERVIEW

period the most important evolutionary phase of the Neuquen Basin started. Thick and widespread successions were deposited during this long period of protracted thermal subsidence and regional back-arc extension. They include a complex series of transgressive-regressive cycles of different magnitude, controlled by the combined effects of changes in subsidence rates, localized uplift and eustatic sea-level oscillations (Cuyo, Lotena and Mendoza groups, Fig. 3).

Late Cretaceous -Cenozoic compression and foreland basin phase Towards the end of the Early Cretaceous changes in the rates of South Atlantic spreading and a reorganization of the Pacific plates, including a decrease in the angle of slab subduction, resulted in the development of a compressional tectonic regime that caused inversion of previous extensional structures (Vergani et al. 1995). At this stage the Neuquen region became a retro-arc foreland basin (Fig. 4C), and significant variations in the size and shape of the basin (Legarreta & Uliana 1991) together with an eastwards migration of the depocentres occurred (Franzese et al 2003). The active depositional systems within the Neuquen Basin were strongly controlled by the compressive regime. Uplift and tectonic inversion in the mountain chain to the west led to the deposition of more than 2000 m of continental deposits in the main depocentres (Rayoso and Neuquen groups, Fig. 3) (Legarreta & Uliana 1991, 1999; Vergani et al. 1995). Towards the end of the Cretaceous continental sedimentation was widespread and the Neuquen Basin merged with other basins to the south (e.g. the San Jorge Basin) to produce a unique giant depocentre (Franzese et al. 2003). In the latest Cretaceous very high global sea levels resulted in the first marine transgression from the Atlantic, with shallow-marine deposits occurring over wide areas of the basin (Barrio 1990). Several thin- and thick-skinned fold and thrust belts developed as a result of the foreland basin phase (Ramos 1999/?) and their position constitutes a major control on the present-day physiography of the Neuquen region (Fig. 2). However, the compressional regime was not a continuous, simple process through time. Zapata & Folguera (this volume) have identified several different stages of tectonic compression and relaxation in the evolution of the Andean Fold and Thrust Belt between the Late Cretaceous and Cenozoic. Moreover, these authors propose that flexural subsidence during

7

tectonic compression was occasionally coeval with the generation of small depocentres associated with intense (arc and retro-arc) volcanic activity (Fig. 2). Ramos & Folguera (this volume) provide a detailed analysis of the main characteristics and evolution of these magmatic-related depocentres.

Chrono- and biostratigraphic framework The development of thick and virtually continuous Jurassic-Early Cretaceous marine successions, together with a complete and varied record of ammonoid, brachiopod, bivalve and microfossil faunas, has contributed to a highly refined biostratigraphy for the basin during this interval. The Jurassic ammonite faunas are one of the most continuous and complete records anywhere in the world. More than 30 ammonite biozones are defined for the Jurassic stages (Leanza 1973, 1981; Riccardi 1983; Riccardi et al. 1990a-c, 1999). The only exception to this almost complete record occurs in the Kimmeridgian, where a major tectonic inversion phase caused a protracted fall in relative sea level and a 7 Ma biostratigraphic gap (Fig. 3) (Riccardi et al. 1999). A similar level of biostratigraphic refinement has been attained for the Early Cretaceous strata (Leanza 1973, 1981; Leanza & Hugo 1977; Aguirre-Urreta & Rawson 1997; AguirreUrreta et al. 1999). The chronostratigraphy of the Berriasian-Barremian interval is further refined using a combination of ammonites, calcareous nannofossils and palynomorphs by Aguirre-Urreta et al. (this volume). The high resolution of the ammonite zones within the basin give a precision of 500 ka for some of the biozones, making the area ideal for basin analysis studies in which time-constrained stratigraphy is essential (e.g. Sagasti this volume; Schwarz & Howell this volume). As the Cretaceous-Tertiary (K/T) boundary can be identified within a marine succession on the basis of microfossil faunas (Casadio et al. 2004), the basin is an ideal site for further research on the causes and effects of K/T global extinctions. In contrast, Mesozoic intervals that are characterized by continental-dominated deposition in the basin (e.g. the Late Trias sic and Late Cretaceous) lack a well-defined stratigraphic framework (Fig. 3). With the exception of a marine Triassic-Early Jurassic succession in the Atuel rift (Riccardi & Iglesia Llanos 1999), the chrono- and biostratigraphic record for the Late Triassic is generally poor. In the case of the Late Cretaceous, much of the record is

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J. A. HO WELL £7 AL.

comprised of continental and arid-marginal marine deposits that include a rich fauna of terrestrial reptiles (Fig. 3), but lack fossils that provide biostratigraphic constrains. The Palaeogene biostratigraphic record is equally poor, although the presence of volcanic horizons related to the arc magmatism provides an important geochronological database (Llambias & Rapela 1989; Ramos 1999a; Jordan et al 2001; Folguera et al. 2004, and references therein).

Jurassic-Cretaceous sequence stratigraphy The majority of the succession that crops out in the Neuquen region was deposited in the Jurassic-Cretaceous post-rift basin. During this period steep subduction of the Pacific plates resulted in negative roll-back and a broad, generally extensional regime in both the arc and backarc settings (Ramos 19996). Within the Neuquen Basin this extension was gentle and was expressed as broad-scale, regional subsidence rather than rifting with active extensional faults at the surface. The depositional systems were marine-dominated and show well-defined records of cyclic sea-level change at different scales. These cycles were a product of the complex interaction of eustatic oscillations, minor extension and thermal subsidence with localized uplift and inversion, and form the focus of sequence stratigraphic studies of the sedimentary record in the basin. In his pioneering study of the stratigraphy in the basin, Groeber (1946) identified two major cycles (Jurasico and Andico), each composed of several transgressive-regressive subcycles. Building on this work, several authors (Gulisano et al. 1984; Mitchum & Uliana 1985; Legarreta & Gulisano 1989; Legarreta & Uliana 1991, 1996, 1999; Legarreta et al. 1993) produced a more detailed breakdown of these cycles and attributed them primarily to eustatic sea-level changes under a regime of thermal subsidence. The dramatic sea-level falls that occurred during the Cretaceous (>100m), such as the sequence boundaries at the base of the Avile and Troncoso members (Fig. 3) in which aeolian deposits overlie offshore marine shales (Veiga et al. 20020; Veiga et al. this volume) were attributed to sea level in the Pacific falling below a sill in the Andean arc that separated the Neuquen Basin from the open ocean. Whilst appealing and an excellent starting point, this interpretation appears to have underrated the importance of intrabasinal and intraarc tectonics. According to Vergani et al.

(1995), Tankard et al. (1995), Pangaro et al. (2002) and Veiga et al. (2002fc), the sag phase of subsidence was frequently disturbed by tectonic reactivations associated with changes in the subduction regime and intraplate reorganization. There are a number of aspects of the basin that make it an excellent case study in sequence stratigraphy. The high-resolution biostratigraphic record provides a framework for study; the high-quality outcrops and the proximity to an abundance of subsurface information provide good data to develop and constrain models, and the geodynamic setting outlined above produced well-developed cycles of relative sea level change. In the early Jurassic the basin had a topography that was inherited from the late Triassic rift phase (Burgess et al. 2000). During the remainder of the Jurassic and Early Cretaceous history the basin had a ramp-style geometry, similar to other retro-arc basins (e.g. the Western Interior Basin of the USA; Edwards et al 2005). The Early Jurassic of the Neuquen Basin provides an excellent study in the significance of basin geometry on sequence and facies architecture. The deep-water turbidite systems of the Los Molles Formation (Burgess et al. 2000) and the shallow-marine tidal deposits of the Lajas Formation (Mcllroy et al. this volume) were strongly influenced by the relict topography inherited from the early rift phase. This topography controlled the distribution of depositional lows, and in the Lajas Formation resulted in the localized amplification of the tidal wave and a thick, highly aggradational succession of tidal deposits. Deposition in the late post-rift ramp setting was characterized by well-developed cycles showing a complete record of lowstand, transgressive and highstand systems tracts. Surfaces that bound these sequences are marked by a sharp basinward shifting of continental-dominated facies. Falling-stage deposits are present in some cases (Veiga et al. this volume), but are typically poorly developed. The transgressive systems tracts are mainly composed of thick offshore deposits (Doyle et al. this volume; Sagasti this volume), even near the basin margins. These deposits commonly show features of restricted marine circulation. The highstand systems tracts are mainly composed of mixed offshore siliciclastics and carbonates that pass upwards into progradational shoreface, deltaic and fluvial deposits (Fig. 3). The Lower Cretaceous succession of the Neuquen Basin includes a number of such examples of rampmargin sequences.

THE NEUQUEN BASIN: AN OVERVIEW

The extreme facies shifts that are associated with the sequence boundaries are attributed to the effects of relative sea-level fall, enhanced and locally overprinted by phases of localized tectonic inversions. Although the basin ward shift in facies is commonly major, the sequence boundaries are typically planar and incised valleys are rare (Schwarz et al. 2005; Schwarz & Howell this volume). The nature of the facies that overlie the sequence boundaries is partially controlled by the degree of connection that was maintained to the protoPacific Ocean. In some cases a complete desiccation of the basin occurred as the connection was severed (e.g. the aeolian deposits of the Troncoso Member: Veiga et al. this volume), in others a limited connection was maintained and the lowstand deposits show evidence of open or restricted marine circulation. Schwarz & Howell examined one of these long-term lowstand wedges, and highlight how tectonic activity and basin physiography conditioned the internal sequence architecture and the relationship between contemporary marine and non-marine depositional systems. The low angle of the ramp margin also favoured rapid landwards migration of shorelines during the transgressions that followed the lowstands. In many cases shallow-marine and offshore deposits directly overlie fluvial and aeolian facies. Stromback et al. (this volume) analysed one of these transgressive events that occurred across the top of a lowstand aeolian sand sea. In this case the transgression was fast enough to preserve at least some of the dune topography with soft-sediment deformation and slumping into the interdune lows, and only localized reworking of the dune tops. Transgressive systems tracts within the postrift fill of the basin are characterized by thick successions of offshore marine deposits that commonly show evidence for restricted water circulation. Within these cyclically stacked black shale and marl successions Doyle et al. (this volume) examined how systematic variation in the Jurassic-Lower Cretaceous ichnological and faunal record may be employed to interpret the firmness of the marine substrate and different levels of oxygenation at the water-sediment interface. Besides, a detailed study of organic facies within transgressive intervals by Tyson et al. (this volume) reveal that Cretaceous anoxic events do not exactly correlate with previously documented global anoxic events. They are interpreted as the result of the combination of a long-term rise in sea level and the development of locally restricted conditions.

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Within low-frequency transgressive cycles, high-frequency subdivisions may be recognized in the Neuquen Basin record. Scasso et al. (this volume) analysed the rhythmic succession of limestones and marls that characterize one of the high-frequency Tithonian highstands, concluding that these offshore cycles are the result of systematic changes in productivity on the sea surface, and supply of terrigenous and nonterrigenous material in suspended plumes. Sagasti (this volume) analysed high-frequency cycles developed during two low-order Valanginian-Barremian transgressive successions. These outer ramp rhythms are interpreted as dilution cycles triggered by orbital climatic changes within the Milankovitch range. Towards the end of the Early Cretaceous the Neuquen Basin started to experience one of its major tectonic changes, passing from the backarc sag phase to the early part of the foreland phase. Veiga et al. (this volume) analysed the sequence stratigraphic architecture and the evolution of the depositional systems through this transition. Some striking differences are depicted from the previous sequence stratigraphic framework, with a well-developed falling-stage systems tract followed by a lowstand episode characterized by complete disconnection from the ocean and without re-establishment of 'normal' marine conditions during the subsequent transgression.

Palaeobiology The biological record of the Neuquen Basin is diverse and continuous, and, in addition to its biostratigraphic significance, it also allows transcendent palaeoecological, taphonomical and palaeobiogeographical studies. As with other studies in the basin, this work exceeds its local significance and contributes to interpretations that are applicable worldwide. The most famous palaeobiological record is that of the Mesozoic reptiles of the Neuquen Basin. So far the most important fossil reptiles of southernmost South America (including Patagonia) all come from the Neuquen Basin. The rich dinosaur fauna has resulted in the definition of many new taxa (Coria & Salgado 1995, 1996; Bonaparte 1996, 1998; Coria 2001; Coria & Calvo 2002, among others), the development of evolutionary models (Wilson & Sereno 1998; Sereno 1999), and the study of faunal assemblages and reptile palaeocommunities (Novas 1997; Leanza et al 2004). Coria & Salgado (this volume) analysed the saurischian dinosaur evolutionary trends and discussed the main causes of intra-Cretaceous extinctions.

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It is not just the dinosaurs of the Neuquen Basin that are outstanding. Marine reptiles are also very well preserved in the Jurassic and Lower Cretaceous successions of the basin, as shown by Gasparini & Fernandez (this volume). In particular, the wonderful record of Late Jurassic marine reptiles has allowed the studies on taphonomy and palaeobiological interactions within an almost isolated marine embayment (Fig. 5). These palaeontological studies have strongly contributed to new palaeobiogeographic panoramas and to the definition of biological connections between different oceanic realms (Gasparini 1996; Gasparini & Fernandez 1997). While the reptile fauna of the basin is dramatic, the Mesozoic invertebrates are equally well preserved and represented. Besides the biostratiographic significance of macro- and microinvertebrate faunas, they have allowed the development of detailed biofacial and taphonomic studies. Lazo et al. (this volume) show the great variability of invertebrate palaeocommunities developed in different subenvironments of the Neuquen marine ramp during the Early Cretaceous.

Despite a number of palynological contributions (cf. Quattrocchio & Sarjeant 1992; Quattrocchio et al. 1996, 2002; Martinez et al. 2005) the mega-palaeofloristic record of the Neuquen Basin is not as well documented. The contribution by Morgans-Bell & Mcllroy (this volume) shows how morphological studies of Jurassic conifers can contribute to palaeoenvironmental and palaeoclimatic interpretations. Perspectives and future work Despite the significant volumes of previous work, including that detailed in this volume, studies of the Neuquen Basin are still in their infancy. Both the outcrops and the subsurface portions of the basin offer significant potential for further work that has global implications. Detailed understanding of the subsurface reservoirs that exist in the Neuquen Embayment is still not in the public domain (if it exists). There are considerable opportunities for further comparison of the producing reservoirs with the outcrops. Outcrop characterization and modelling, compared and contrasted to oil-field production data from the same intervals less than

Fig. 5. Reconstruction of the Tithonian marine herpetofauna of the Neuquen Basin (original drawing by J. Gonzalez, courtesy of Dr Z.B. de Gasparini).

THE NEUQUEN BASIN: AN OVERVIEW 50 km apart, provides potential for numerous studies. As does linking the well log and seismic expression of the intervals to their outcrop expression. The subsurface data also hold the key to many of the unsolved palaeogeographic problems, and the potential for high-quality, unweathered biostratigraphic data from cores is far reaching. When compared with other parts of the world with comparable outcrop quality, the outcrops of the Neuquen Basin have received little attention. In the future, further studies will be undertaken to improve our understanding of facies and sequence stratigraphy. There is considerable scope for inversion and forward modelling of the observed stratigraphic architecture, and such work will be central to understanding the details of the driving mechanisms behind the dramatic sea-level falls and rapid flooding surfaces that have been documented, and the timing and duration of the lowstands. There is also considerable scope for the development of depositional models and high-resolution sequence stratigraphic schemes for the synrift and foreland stages of the basin history. Such studies will be highly dependent on the construction of more complete chrono stratigraphic and biostratigraphic framework for these stages. Whilst the stratigraphic scheme for much of the basin history is very good, further attention must be paid to more absolute dating of the volcanic and volcaniclastic rocks. This will result in a refinement of the current biostratigraphic schemes for the Jurassic and Cretaceous, and an improved understanding of the Triassic and Cenozoic histories. Further improvements of the stratigraphy of the basin will also arise from much greater integration of the existing and future subsurface data. Much of our existing knowledge of the basin fill is taken from the outcrops towards the SE and NE (passive, cratonic) margins of the basin. The geometry and physiography of the western (active) margin of the basin are far less well understood. In the near future, studies on the Jurassic and Cretaceous sedimentary record close to the magmatic arc will be required to define the main sedimentary processes, to validate sequence stratigraphic schemes and to locate the pathways across the magmatic arc that allowed connection of the Neuquen Basin with the proto-Pacific Ocean. Excellent outcrops, copious subsurface data, a world class palaeontological record and a unique structural setting combine to make the Neuquen Basin a unique case study in basin evolution and fill. This Special Publication represents the state of current understanding and hopefully

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highlights the enormous potential for future study.

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GROEBER, P., STIPANICIC, P.N. & MINGRAMM, A. 1953. Jurasico. Geografia de la Republica Argentina, 2, 143-347. GULISANO, C.A., GUTIERREZ PLEIMLING, A.R. & DIGREGORIO, R.E. 1984. Analisis estratigrafico del intervalo Tithoniano-Valanginiano (Formaciones Vaca Muerta, Quintuco y Mulichinco) en el suroeste de la provincia de Neuquen. IX Congreso Geologico Argentina, S. C. de Bariloche, 1, 221-235. HERRERO-DUCLOUX, A. 1946. Contribution al conoci miento del Neuquen extrandino. Boletin de Informaciones Petroleras, 22, 245-261. INTROCASO, A., PACING, M.C. & FRAGA, H. 1992. Gravity, isostasy and Andean crustal shortening between latitudes 30° and 35°S. Tectonophysics, 205,31-48. JORDAN, T., BURNS, W., VEIGA, R., PANGARO, F., COPELAND, P., KELLEY, S. & MPODOZIS, C. 2001. Extension and basin formation in the Southern Andes caused by increased convergence rate: A Mid-Cenozoic trigger for the Andes. Tectonics, 20, 308-324. LLAMBIAS, EJ. & RAPELA, C.W. 1989. Las volcanitas de Collipilli, Neuquen (37°S) y su relation con otras unidades paleogenas de la cordillera. Asociacion Geologica Argentina Revista, 44, 224—236. LEANZA, H.A. 1973. Estudio sobre los cambios faciales de los estratos limitrofes jurasico-cretacicos entre Loncopue y Piciin Leufu, Provincia del Neuquen, Republica Argentina. Revista de la Sociedad Geologica Argentina, 28, 97-132. LEANZA, H.A. 1981. Faunas de ammonites del Jurasico superior y del Cretacico inferior de America del Sur, con especial consideration de la Argentina. In: VOLKHEIMER, W. & MussACHio, E. (eds) Cuencas Sedimentarias del Jurasico y Cretacico de America del Sur. Museo Argentine de Ciencias Naturales, Buenos Aires, Argentina, 2, 559-597. LEANZA, H.A. & HUGO, C.A. 1977. Sucesion de amonites y edad de la Formacion Vaca Muerta y sincronicas entre los paralelos 35° y 40° L.S. Cuenca Neuquina-Mendocina. Revista de la Asociacion Geologica Argentina, 32, 248-264. LEANZA, H.A., MARCHESE, H. & RIGGI, J. 1977. Estratigraffa del Grupo Mendoza con especial referencia a la Formacion Vaca Muerta entre los paralelos 35° y 40° L.S., Cuenca Neuquina-Mendocina. Revista de la Asociacion Geologica Argentina, 32, 190-208. LEANZA, H.A., APESTEGUIA, S., NOVAS, F. & DE L. FUENTE, M. 2004. Cretaceous terrestrial beds from the Neuquen Basin (Argentina) and their tetrapod assemblages. Cretaceous Research, 25, 61-87. LEGARRETA, L. & GULISANO, C.A. 1989. Analisis estratigrafico secuencial de la Cuenca Neuquina (Triasico superior-Terciario inferior, Argentina). In: CHEBLI, G. & SPALLETTI, L. (eds) Cuencas Sedimentarias Argentinas. Serie Correlation Geologica, Universidad Nacional de Tucuman, 6, 221-243. LEGARRETA, L. & ULIANA, M.A. 1991. JurassicCretaceous marine oscillations and geometry of

THE NEUQUEN BASIN: AN OVERVIEW back arc basin fill, Central Argentine Andes. In: MACDONALD, D.I.M. (ed.) Sedimentation, Tectonics and Eustasy - Sea-level Changes at Active Margins. International Association of Sedimentologists, Special Publications, 12, 429-450. LEGARRETA, L. & ULIANA, M. 1996. The Jurassic succession in west-central Argentina: stratal patterns, sequences and paleogeographic evolution. Palaeogeography, Palaeoclimatology, Palaeoecology, 120, 303-330. LEGARRETA, L. & ULIANA, M. 1999. El Jurasico y Cretacico de la Cordillera Principal y la Cuenca Neuquina. 1. Facies Sedimentarias. In: CAMINOS, R. (ed.) Geologia Argentina. Instituto de Geologia y Recursos Minerales, Anales, 29, 399-416. LEGARRETA, L., GULISANO, C.A. & ULIANA, M.A. 1993. Las secuencias sedimentarias jurasico-cretacicas. In: Relatorio Geologia y Recursos Minerales de Mendoza. XII Congreso Geologico Argentino y II Congreso de Exploracion de Hidrocarburos, Mendoza, 1(9), 87-114. MACDONALD, D., GOMEZ-PEREZ, I. ET AL. 2003. Mesozoic break-up of SW Gondwana: Implications for South Atlantic regional hydrocarbon potential. Marine and Petroleum Geology, 20, 287-308. MANCEDA, R. & FIGUEROA, D. 1995. Inversion of the Mesozoic Neuquen rift in the Malargiie fold and thrust belt, Mendoza, Argentina. In: TANKARD, A.J., SUAREZ SORUCO, R. & WELSINK, HJ. (eds) Petroleum Basins of South America. AAPG Memoirs, 62, 369-382. MARTINEZ, M.A., QUATTROCCHIO, M.E. & PRAMPARO, M.B. 2005. Analisis palinologico de la Formacion Los Molles, Grupo Cuyo, Jurasico medio de la cuenca Neuquina, Argentina. Ameghiniana, 42, 67-92. MITCHUM, R.M. & ULIANA, M.A. 1985. Seismic stratigraphy of carbonate depositional sequences, Upper Jurassic-Lower Cretaceous, Neuquen Basin, Argentina. In: BERO, B.R. & WOOLVERTON, D.G. (eds) Seismic Stratigraphy: An Integrated Approach to Hydrocarbon Exploration. AAPG Memoirs, 39, 255-274. NOVAS, F.E. 1997. South American dinosaurs. In: CURRIE, P. & PADIAN, K. (eds) Encyclopedia of Dinosaurs. Academic Press, San Diego, CA, 678-689. PANGARO, F., VEIGA, R. & VERGANI, G. 2002. Evolucion tecto-sedimentaria del area de Cerro Bandera, Cuenca Neuquina, Argentina. V Congreso Argentino de Hidrocarburos, Mar del Plata (electronic format), IAPG, Buenos Aires, Argentina. PLOSZKIEWICZ, J.V., ORCHUELA, LA., VAILLARD, J.C. & VINES, R.F. 1984. Compresion y desplazamiento lateral en la zona de falla de Huincul, estructuras asociadas, provincia del Neuquen. IX Congreso Geologico Argentino, San Carlos de Bariloche, 2, 163-169. QUATTROCCHIO, M.E. & SARJEANT, W.A.S. 1992. Dinoflagellate cysts and acritarchs from the Middle and Upper Jurassic of the Neuquen Basin, Argentina. Revista Espanola de Micropaleontolo£ifl,24,67-118.

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QUATTROCCHIO, M.E., MARTINEZ, M.A., GARCIA, V.M. & ZAVALA, C.A. 2002. Palinoestratigrafia del Tithoniano-Hauteriviano del centro-oeste de la Cuenca Neuquina, Argentina. In: 8th Congreso Argentino de Paleontologia y Bioestratigrafia, Resumenes. Corrientes, Argentina, 75-76. QUATTROCCHIO, M.E., SARJEANT, W.A.S. & VOLKHEIMER, W. 1996. Marine and terrestrial Jurassic microfloras of the Neuquen Basin (Argentina): palynological zonation. GeoResearch Forum, 1-2, 167-178. RAMOS, V.A. 1978. Estructura. In: Relatorio Geologi y Recursos Naturales del Neuquen. VII Congreso Geologico Argentino, Neuquen, 99-125. RAMOS, V.A. \999a. Los depositos sinorogenicos terciarios de la region andina. In: CAMINOS, R. (ed.) Geologia Argentina. Instituto de Geologia y Recursos Minerales, Anales, 29, 651-682. RAMOS, V.A. 19996. Evolucion Tectonica de la Argentina. In: CAMINOS, R. (ed.) Geologia Argentina. Instituto de Geologia y Recursos Minerales, Anales, 29, 715-759. RICCARDI, A.C. 1983. The Jurassic of Argentina and Chile. In: MOULLADE, M. & NAIRN, A.E. (eds The Phanerozoic Geology of the World, II. The Mesozoic. Elsevier, Amsterdam, 201—263. RICCARDI, A.C. & IGLESIA LLANOS, M.P. 1999. Primer hallazgo de amonites en el Triasico de la Argentina. Revista de la Asociacion Geologica Argentina, 54, 298-300. RICCARDI, A.C., WESTERMANN, G.E.G. & LEVY, R. 1971. The Lower Cretaceous Ammonitina Olcostephanus, Leopoldia and Favrella from west-central Argentina. Palaeontographica, 136, 83-121. RICCARDI, A.C., DAMBORENEA, S.E. & MANCENIDO, M.O. 1990(2. Lower Jurassic of South America and Antarctic Peninsula. In: WESTERMANN, G.E.G. & RICCARDI, A.C. (eds) Jurassic Taxa Ranges and Correlation Charts for the CircumPacific. Newsletters on Stratigraphy, 21(2), 75-103. RICCARDI, A.C., DAMBORENEA, S.E. & WESTERMANN, G.E.G. 19906. Middle Jurassic of South America and Antarctic Peninsula. In: WESTERMANN, G.E.G. & RICCARDI, A.C. (eds) Jurassic Taxa Ranges and Correlation Charts for the Circum-Pacific. Newsletters on Stratigraphy, 21(2), 105-128. RICCARDI, A.C., LEANZA, H.A. & VOLKHEIMER, W. 1990c. Upper Jurassic of South America and Antarctic Peninsula. In: WESTERMANN, G.E.G. & RICCARDI, A.C. (eds) Jurassic Taxa Ranges and Correlation Charts for the Circum-Pacific. Newsletters on Stratigraphy, 21 (2), 129-147. RICCARDI, A.C., DAMBORENEA, S.E. & MANCENIDO, M.O. 1999. El Jurasico y Cretacico de la Cordiller Principal y la Cuenca Neuquina. 3. Bioestratigrafia. In: CAMINOS, R. (ed.) Geologia Argentina. Institut de Geologia y Recursos Minerales, Anales, 29, 419-432. SCHWARZ, E., SPALLETTI, L.A. & HOWELL, J.A. 2005. Sedimentary response to a tectonically-induced sea-level fall in a shallow back-arc basin: the

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Mulichinco Formation (Lower Cretaceous), Neuquen Basin, Argentina. Sedimentology, in press. SERENO, P.C. 1999. The evolution of dinosaurs. Science, 284, 2137-2147. SPALLETTI, L., FRANZESE, J., MATHEOS, S., & SCHWARZ, E. 2000. Sequence stratigraphy of a tidally-dominated carbonate-siliciclastic ramp; the Tithonian of the southern Neuquen Basin, Argentina. Journal of the Geological Society, London, 157, 433-446. TANKARD, A.J., ULIANA, M.A. ETAL. 1995. Structural and tectonic controls of basin evolution in southwestern Gondwana during the Phanerozoic. In: TANKARD, A.J., SUAREZ SORUCO, R. & WELSINK, HJ. (eds) Petroleum Basins of South America. AAPG Memoirs, 62, 5-52. VALENTE, S. 1999. Modelo deposicional para las areniscas inferiores de la Formacion Mulichinco, Dorso de los Chihuidos, Neuquen. IV Congreso de Exploration y Desarrollo de Hidrocarburos, Mar del Plata, 2, 749-771. VEIGA, G.D., SPALLETTI, L.A. & FLINT, S. 2002o. Aeolian/fluvial interactions and high resolution sequence stratigraphy of a non-marine lowstand wedge: The A vile Member of the Argio Formation (Lower Cretaceous) in central Neuquen Basin, Argentina. Sedimentology, 49, 1001-1019. VEIGA, R., PANGARO, F. & FERNANDEZ, M. 20026. Modelado bidimensional y migracion de hidrocarburos en el ambito occidental de la Dorsal de

Huincul, Cuenca Neuquina-Argentina. V Congreso Argentino de Hidrocarburos, Mar del Plata (electronic format), IAPG, Buenos Aires, Argentina. VERGANI, G.D., TANKARD, A.J., BELOTTI, HJ. & WELSINK, HJ. 1995. Tectonic evolution and paleogeography of the Neuquen Basin, Argentina. In: TANKARD, A.J., SUAREZ SORUCO, R. & WELSINK, HJ. (eds) Petroleum Basins of South America. AAPG Memoirs, 62, 383-402. VERGANI, G.D., SELVA, G. & BOGGETTI, D.A. 2002. Estratigrafia y modelo de facies del Miembro Troncoso Inferior, Formacion Huitrin (Aptiano), en el noroeste de la Cuenca Neuquina, Argentina. In: CINGOLANI, C.A., CABALERI, N., LINARES, E., LOPEZ DE LUCHI, M.G., OSTERA, H.A. & PANARELLO, H.O. (eds) XV Congreso Geologico Argentino, El Calafate, 1, 613-618. WEAVER, C.E. 1931. Paleontology of the Jurassic and Cretaceous of West Central Argentina. Memoir of the University of Washington, 1, 1 -469. WILSON, J.A. & SERENO, P.C. 1998. Early Evolution and Higher-Level Phylogeny of Sauropod Dinosaurs. Society of Vertebrate Paleontology, Memoirs, 5, 1-68. YRIGOYEN, M.R. 1991. Hydrocarbon resources from Argentina. In: World Petroleum Congress, Buenos Aires. Petrotecnia, 13, Special Issue, 3854.

Tectonic evolution of the Andes of Neuquen: constraints derived from the magmatic arc and foreland deformation VICTOR A. RAMOS & ANDRES FOLGUERA Laboratorio de Tectonica Andina, Facultad de Ciencias Exactas y Naturales, Universidad de Buenos Aires, Ciudad Universitaria, 1428 Buenos Aires, Argentina (e-mail: [email protected]) Abstract: The Andes of the Neuquen region (36°-38°S latitude) of the Central Andes have distinctive characteristics that result from the alternation of periods of generalized extension followed by periods of compression. As a result of these processes the Loncopue trough is a unique long depression at the foothills parallel to the Principal Cordillera that consists of a complex half-graben system produced during Oligocene times and extensionally reactivated in the Pliocene-Pleistocene. Its northern sector represents the present contractional orogenic front. The nature and volume of arc-related igneous rocks, the location of the volcanic fronts, expansions and retreats of the magmatism, and the associated igneous activity in the foreland, together with the analyses of the superimposed structural styles, permit the constraint of the alternating tectonic regimes. On these bases, different stages from Jurassic to Present are correlated with changes in the geometry of the Benioff zone through time. Periods of subduction-zone steepening are associated with large volumes of poorly evolved magmas and generalized extension, while shallowing of the subduction zone is linked to foreland migration of more evolved magmas associated with contraction and uplift in the Principal Cordillera. The injection of hot asthenospheric material from the subcontinental mantle into the asthenospheric wedge during steepening of the subduction zone produced melting and poorly evolved magmas in an extensional setting. These periods are linked to oceanic plate reorganizations in the late Oligocene and in the early Pliocene.

The tectonic evolution of the sub-Andean Neuquen Basin is a consequence of the interaction of different processes along the continental margin. The geological history of the Andes in the Neuquen region is somewhat different to the rest of the Central Andes. Most of the fault segments between the Guayaquil (4°S latitude) and the Penas (46°30/S latitude) gulfs have active orogenic fronts that have been under contraction since the late Cenozoic (Allmendinger et al 1997; Ramos 1999; Jaillard et al 2000; Ramos & Aleman 2000). As a consequence of subduction erosion and changes in the geometry of the Wadati-Benioff zone the magmatic arcs of the Central Andes have shifted towards the foreland during the Late Cretaceous-late Cenozoic Andean cycle (Kay et al. 1987; Mpodozis & Ramos 1989; Ramos et al. 1991; Kay 2002). Secondly, with the exception of the Pampean flat-slab segment, the orogenic fronts segments are located between the thrust front and the undeformed foreland (Jordan et al. 1983; Ramos et al. 2002). The foothills of

these regions concentrate most of the intraplate shallow seismicity, and earthquake epicentres are related to the wedge top of the fold-and-thrust belts that coincide with the active contraction of the foreland system (DeCelles & Gilest 1996). The orogenic front in the study region is now contracting the Plio-Pleistocene arc, westward from the Neogene fold and thrust belt that is currently inactive. The Neuquen Andes record an oscillatory behaviour since the Jurassic, with the shifting and expansion of the location of arc magmatism of the order of a few tens of kilometres. This is in contrast with the evolution of the other segments of the Central Andes that record arc migrations to the foreland of 400-750 km from the trench (Fig. 1). Consequently, the Neuquen Andes have, at present, an extensional depression between 36°30/ and 39°00'S, known as the Loncopue Graben (Ramos 1977), which is absent from the segments. The Loncopue Graben is located parallel to the cordilleran axis, between the foothills and the foreland region, and the present

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 15-35. 0305-8719/05/$15.000...© The Geological Society of London 2005.

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Fig. 1. Location of the Neuquen Andes within the Central Andes showing the maximum expansion of the volcanic front towards the foreland in the different magmatic arcs (based on Mpodozis & Ramos 1989).

morphology was developed during late Cenozoic times. The present erogenic front is located in the inner retro-arc sector, west of a fossil fold and thrust belt developed during Late Cretaceous and Miocene times. The objective of this paper is to describe the geology along the axis of the Cordillera and the main characteristics of the Loncopue Graben in order to reconstruct the evolution of the Neuquen Andes and the geological history of the adjacent Neuquen Embayment. There is a close relationship between the age of foreland migration of the magmatic arc and deformation, and the age of extensional collapse and large volumes of igneous activity during the period of arc retreat. The study area comprises the main Andes and the foothills between 36° and 39°S. Most of the information is derived from extensive field work and mapping along the foothills and in the inner region of the cordillera (Folguera & Ramos 2000; Folguera et al 20030, b, 2004) (Fig. 2). Data from the forearc and the western

slope are based on the observations of Suarez & Emparan (1997), Melnick et al. (2002, 2005) and Radic et al. (2002). Present tectonic setting An outstanding feature of the Andes within the study area is the Loncopue Graben (also called the Loncopue trough), a morphological depression 300 km long and about 30-40 km wide (Fig. 3). The graben is located between the eastern foothills of the Andes and a fossil antithetic belt (sensu Roeder 1973) called the Agrio Fold and Thrust Belt (Ramos 1977). The Agrio Fold and Thrust Belt has a long and complex long history, and the main contractional deformation ended in the late Miocene (see Zapata & Folguera 2005). The Loncopue trough south of Loncopue is presently bounded to the east by Quaternary normal faults. Triangle facets and recent scarps indicate neotectonic activity along these faults (see Garcia Morabito 2004). These neotectonic features may be

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Fig. 2. Topographic map of the Neuquen Andes showing the location of the present volcanic arc and the extension of the Loncopue depression in the foothills (based on IUGS digital elevation model). Arrows indicate the eastern tectonic boundary of the Loncopue trough.

inherited from an older normal fault developed during the Oligocene that consists of a W-dipping extensional fault that is seen in seismic lines along the eastern margin in the northern sector of the depression (Jordan et al. 2001). A large part of the depression is covered by Quaternary alkaline basalts, as first described by Munoz & Stern (1985). Several lava flows, tens of metres thick, have flowed to the east. Many monogenic pyroclastic cones of basaltic composition are spread over the area. Scarce geochronological data, mainly from the southern end of the depression at about 39°S, indicate ages between 2.30 ± 0.3 and 0.47 + 0.2 Ma (K-Ar whole rock: Linares & Gonzalez 1990). Similar ages of 0.130 ± 0.02 and 0.167 ± 0.005 Ma by Ar-Ar have been reported by Rabassa et al. (1987) in basaltic lavas further south along the same structure. Other Quaternary ages between 1.6 ± 0.2 and 0.9 ± 0.3 Ma (by K-Ar whole rock) have been obtained by Munoz & Stern (1985, 1988) around Paso Pino Hachado.

Interbedded basaltic flows or overlying glacial deposits have also been reported along the Loncopue trough (Folguera et al. 2003&). The poorly evolved alkaline magmatism that characterizes the thick lava flows, the low 87 Sr/86Sr initial ratios near 0.7040, similar to the main erogenic arc (Munoz Bravo et al. 1989), as well as the abundance of monogenic small volcanoes, together with scattered evidence of normal faults as depicted by Folguera et al. (2004), indicate an extensional regime in the retro-arc during Pleistocene times. The Principal Cordillera at these latitudes is bounded to the west by a Holocene volcanic front, and is located 250 km east of and parallel to the trench with a NNE trend. There are a few isolated volcanoes along the axis, such as Llaima, Callaqui, Copahue, Antuco and Chilian (Fig. 3). Geochemical composition and 87 Sr/86Sr ratios ranging from 0.7038 to 0.7041, independent of SiO2 content in the main orogenic arc, indicate crystal fractionation without

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Fig. 3. Location of the Loncopue trough with the main tectonic elements. Note the position and trend of the present volcanic front in comparison with the Quaternary retro-arc basalts of the Loncopue trough. L.O.F.Z., Liquine-Ofqui Fault Zone.

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significant contamination by the crust as the dominant process in an extensional regime similar to the retro-arc (Munoz & Stern 1988). Geophysical data, mainly gravity and preliminary seismological surveys, indicate an unusually thin crustal root beneath the Neuquen Andes. Recent data on receiver function beneath Neuquen at 39°S indicate an abnormally thin crust beneath the Loncopue trough (Kind

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et al 2001). The interpretation of the crustal structure (Fig. 4) is based on the gravity surveys of Couch et al. (1981), Pacino (1993) and Martinez et al (1997), as well as on the receiver function data. The gravity data feature a small crustal root, less than 42-43 km deep. This poses important constraints on the structural style of the Agrio Fold and Thrust Belt, as it implies that crustal shortening cannot exceed

Fig. 4. Location of the earthquake epicentres and crustal section of the Neuquen Andes showing the present Benioff zone (based on Kind et al. 2001; Bohm et al. 2002; Ramos et al. 2004).

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44 km at 37°S and 20 km at 39°S (Martinez et al 1997) if it is assumed that arid conditions prevailed since the late Cretaceous and no significant erosion took place. This is in accordance with the Moho depth observed beneath the main Andes from broadband seismology by Kind et al (2001). Preliminary receiver function data, obtained by broadband teleseismic stations along a west-east profile between the trench and the Neuquen Embayment just a few kilometres north of 39°S, confirm a maximum thickness of between 40 and 45 km beneath the Neuquen Andes (Kind et al. 2001). Significantly, the minimum crustal thickness is reported west of 70.2°W longitude, below the Loncopue trough. Although these data are preliminary, a shallow Moho is reported at less than 30 km depth, considerably less than the 40 km recorded in the plains of the Neuquen Embayment. If these data are confirmed it would be a clear indication of significant crustal thinning developed in the retro-arc during an extensional regime that began in the Oligocene, was interrupted in the late Miocene and lasted until the Pleistocene. Seismological data reported by INPRES indicate different behaviour within the foothills north and south of 37°30/S (Fig. 4). The northern area records an active orogenic front with more frequent intraplate earthquakes that coincides with the northern sector of the Loncopue trough where active compressional neotectonics has been recently described (Folguera et al. 20030, 2004). The southern area is far less active in comparison and there is no evidence of compressional neotectonic features. Previous magmatic arcs and intra-arc basins In order to understand the present tectonic framework, it is important to analyse the magmatic history of the volcanic arc through time. This includes the location of the volcanic front, the characteristics of the volcanic products, the basin formation and the subsequent tectonic regime. In the following section several different magmatic episodes are described as well as the resulting tectonic regime. These episodes include: the Jurassic-Early Cretaceous; the Late Cretaceous-Palaeogene; the Oligoceneearly Miocene; the middle-late Miocene; and the Pliocene-Pleistocene. The outstanding characteristic of these episodes is the oscillatory nature of the migration and expansion of magmatic activity, and a somewhat stationary volcanic front (Ramos 1988; Mpodozis & Ramos

1989). The subduction complex is preserved along the Pacific margin at these latitudes and it has no evidence of erosion. The lack of important subduction erosion as depicted further north by Stern (1991) and Kay (2002), rules out this mechanism as a cause for the migration and expansion of the magmatic activity toward the foreland.

Jurassic-Early Cretaceous arc and intra-arc basin Volcanic and plutonic rocks of this age are widely preserved along the axis of the cordillera north of 36°S. However, to the south of this latitude, scarce outcrops are partially exposed beneath thick covers of Cenozoic volcanic rocks. For example, west of Cordillera del Viento, between Bella Vista and Nahueve (c. 37°S), there are volcanic domes and necks of andesitic-dacitic composition that have been dated at 167.7 ± 8.2 Ma (K-Ar whole rock) by Rovere (1998). These volcanic rocks have a typical calc-alkaline composition (56.04% SiO2; 1.16% K2O) and are correlated with the arc volcanism developed further north. Jurassic volcanic rocks have also been described near Lonquimay, at 38°30/S, in the lower and upper members of the Nacientes del Biobio Formation by De la Cruz & Suarez (1997). The age of these volcanic sequences is constrained between the Lower and Upper Jurassic based on the interbedded sedimentary facies with abundant ammonites. The tholeiitic basalts of the lower member have been assigned to a magmatic arc developed in an extensional regime in a wide intra-arc basin that reached the Argentine side of the Andes (Ramos 1999). The upper member has been correlated with Kimmeridgian volcanic rocks well developed north of 36°S. The batholith of the Principal Cordillera, from Temuco (c. 38°30'S) to the south, consists of granitoids ranging in age from 176 to 164 Ma (Rb-Sr isochrones: Munizaga et al. 1988; Niemeyer & Mufioz 1983). Similar granitoids of Cretaceous age, emplaced at these latitudes in the Principal Cordillera, yielded ages of 94 + 2 Ma, showing the wide distribution of the Mesozoic granitic rocks along the axis of the cordillera. Based on the general characteristics of the Jurassic and Early Cretaceous rocks where they are well exposed in the adjacent areas, an extensional regime was suggested for these igneous rocks between 35° and 39°S by Munoz (1984). Similar conclusions were obtained by De la

TECTONIC EVOLUTION OF THE NEUQUEN ANDES

Cruz & Suarez (1997) who proposed a volcanic arc front west of the Principal Cordillera. However, the scarcity of exposures precludes a precise palaeogeographic reconstruction. It is generally accepted that they were emplaced during an important period of negative trench roll-back velocity when extensional conditions were widespread in the arc and retro-arc regions (Ramos 1999). Late Cretaceous-Palaeogene arc There was a striking change in the distribution of volcanic and plutonic rocks in the Late Cretaceous when an important expansion of magmatism to the foreland occurred. Within the Principal Cordillera there are many granitoids of Late Cretaceous age that sit within a main batholith which forms the roots of the magmatic arc. K-Ar ages between 36° and 38°S range from 90.36 ± 3.63, 85.4 ± 5.2 and 83.9 ± 3.8 to 76.5 ± 1.8 and 64.0 ± 1.9 Ma (Munizaga et al 1985). There was a migration of the magmatism to the foreland soon after the emplacement of the batholith that coincides with a magmatic lull of Eocene volcanic rocks along the Chilean side of the Andes at 36°-39°S (Lopez Escobar & Vergara 1997). The important Late Cretaceous change along the Pacific margin of South America coincides with a well-documented adjustment of the tectonic regime associated with the final breakaway of the South American Plate from the African Plate and the beginning of the drift stage (Somoza 1995). This transition caused a change from a negative to a positive roll-back velocity, marking the beginning of the contraction along the continental margin (Daly 1989). Volcanic rocks of Late Cretaceous-Eocene age are widely developed between 36° 30' and 38°30/S (Fig. 5). These rocks have been described by Llambias & Rapela (1987, 1989) as part of a calc-alkaline suite of volcanic domes, dykes and lavas that range in composition from basaltic to andesitic. New geochronological data and geochemical analysis of this belt have recently been presented by Franchini et al. (2003). The oldest rocks belong to the Campana Mahuida igneous complex, with an age of 74.2 ± 1.4 Ma, interpreted as a porphyry copper system by Sillitoe (1977). However, recent K-Ar dating indicates a younger age of 60.7 + 1.9 Ma (Franchini et al. 2003). These intrusive rocks have La-Yb ratios of between 10 and 30, typical of a thick crust as seen further north in the Quaternary magmatic arc (Kay et al. 1991; Franchini et al. 2003).

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The granitoids of the Cerro Nevazon area are stocks, sills and dykes of intermediate composition. They range from gabbro to granodiorite, diorite being the dominant facies. They are metaluminous rocks, with a normal potassium content (0.6-2.3%). The La-Yb ratio ranges from 5 to 15, and probably indicates a thinner crust in the northern sector. The age of these rocks is constrained between 56.0 ±1.7 and 59.6 ± 10.6 Ma (K-Ar in hornblende) (Franchini et al. 2003). Further south, in the Quebrada Mala and Cerro Pelan east of Cordillera del Viento, andesitic sills yielded a K-Ar age of 71.5 ± 5 Ma (Llambias & Rapela 1989). Another magmatic system, known as Los Maitenes-El Salvaje was emplaced in the southern end of the Cordillera del Viento. There, a tonalitic stock has an age of 64.7 ± 3.2 Ma (Domfnguez et al. 1984). Similar Paleocene ages were obtained in a stock near Varvarco, NW of Cordillera del Viento (Fig. 5). The Varvarco tonalite yielded an age of 64.7 ± 3.0 Ma (K-Ar whole rock) (JICA 2000). The regions of Caicayen and Collipilli in the central part of this belt preserve a series of intrusive domes and volcanic rocks. They have been assigned to the Collipilli Formation, which is predominantly composed of concordant intrusive bodies like sills and laccoliths. They range in composition from hornblende andesites to diorites and quartz-diorites. There are also small dacitic intrusives. They are calc-alkaline, with normal potassium, and have been interpreted as typical magmatic arc rocks by Llambias & Rapela (1989). A microdiorite from Las Mellizas yielded an age of 49.9 + 3.2 Ma, while a laccolith at Cerro del Diablo had an age of 48.4 + 2.4 Ma; a similar age has been obtained in Cerro Caicayen with 44.7 ± 2.2 Ma (K-Ar whole rock) (Llambias & Rapela 1989). For a sill at Cerro Mayal, Cobbold & Rossello (2003) have recently obtained a late Eocene age (39.7+ 0.2 Ma) by Ar/Ar on whole rock. New Ar/Ar ages from a coarse-grained granodioritic plutonic unit, north of the town of Varvarco in the Cordillera del Viento, were interpreted as cooling ages by Kay (2001). The age of 69.09 ± 0.13 Ma (Ar/Ar in biotite) indicates at least 3 km of uplift during the late Cretaceous, with an extra 3 km uplift prior to the deposition of the Serie Andesitica lavas (Kay 2001). These volcanic rocks, known as the Cayanta Formation, consist of hornblende andesite lava flows, volcanic breccias and volcanic agglomerates widely exposed west of Andacollo town (Llambias & Rapela 1989). New Ar/Ar ages in hornblende dates these rocks at 56.9 + 1.1, 56.5 + 0.6 and 50.3 + 0.6 Ma (Jordan et al. 2001) within the

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Fig. 5. Migration of the location of plutonic and volcanic arc rocks, with the magmatic front during Late CretaceousPalaeogene times indicated (based on Munizaga et al. 1985; Llambias & Rapela 1989; Jordan et al. 2001; Franchini et al 2003). Note that in latest Cretaceous time the magmatic front was east of the Loncopue trough.

same rank of the diorites located east of Cordillera del Viento. As a whole, these rocks indicate that a belt of normal arc characteristics developed to the of east of the Jurassic-Early Cretaceous arc emplaced in a normal-thick crust. The change in composition and nature of the plutono-volcanic arc indicates a thickening episode produced during the Late Cretaceous, with an estimated uplift in the order of 6 km in the northern segment as a result of eastward migration of the magmatic arc toward the foreland. Previous authors have proposed an important transpression during the late Eocene in the northern part of the Neuquen Basin (e.g. Cobbold et al. 1999). This interpretion is based on the orientation of subvertical bitumen veins emplaced mainly in the Jurassic and Cretaceous

rocks (Cobbold et al. 1999). The lack of important Palaeogene synorogenic sequences south of 37°30r S may indicate either that the deformation was concentrated in the inner sector of the Andes or that it was milder than in the northern sector. New findings of growth strata along the flank of the Cortaderas Fault have been interpreted as evidence of Oligocene transpression (Cobbold & Rossello 2002, 2003), but may correspond to a Miocene reactivation depending on the age assigned to this sequence. Oligocene-early Miocene arc and intra-arc basin During this period the volcanic activity between 37° and 41°S underwent further important change in location and character. The Arauco

TECTONIC EVOLUTION OF THE NEUQUEN ANDES

and Valdivia coal basins show evidence that an extensional regime controlled the sedimentation of continental deposits in half-graben systems with a NNE trend along the Coastal Cordillera and the Central Valley during the Oligocene early Miocene (Cisternas & Frutos 1994). These basins received abundant volcaniclastic and pyroclastic deposits from the Principal Cordillera. Andesites and dacites from the western flank of the Andes between 37.5° and 39°S are geochemically and isotopically similar to Quaternary andesites and dacites of the Nevados de Chilian Volcanic Group (36.8°S), which are some of the most primitive andesites and dacites of the Southern Volcanic Zone of the Andes (Lopez Escobar & Vergara 1997). These volcanic sequences were deposited in halfgraben systems (Radic et al 2002), with thicknesses up to 1500m observed on subsurface data. These rocks with flat rare earth element (REE) patterns and La-Yb ratios close to 1 are interpreted to have been erupted during an extensional regime (Vergara et al. 1991 a, b). The lower part of the sequence was included in the Cura Mallin Formation by Suarez & Emparan (1997). It consists of lacustrine and fluvial deposits, interfingered with volcanic and volcaniclastic rocks. The Cura Mallin Basin was described by Radic et al. (2002) as an extensional basin with two depocentres of different polarities (Fig. 6). The northern sub-basin has dominant west-dipping normal faults with a NNE trend, and a 2800 m-thick succession along the eastern border of the Loncopue trough at 37°S. A seismic line presented by Jordan et al. (2001) shows a normal west-dipping fault bounding the eastern margin of the basin west of Andacollo. The volcanic, alluvial and lacustrine deposits of the Cura Mallin Formation were deposited between 24.6 and 22.8 Ma (Ar/Ar: Jordan et al. 2001), and are overlain by the Trapa-Trapa Formation, a thick andesitic pile deposited between 18.2 and 14.7 Ma. The southern sub-basin with a maximum thickness of 2400 m has opposite polarity with a NE trend where the volcanic, alluvial and lacustrine deposits of Cura Mallin Formation were formed between 19.9 and 10.7 Ma. This southern succession is covered by the Mitrauquen Formation, another thick pile of andesites and dacites deposited between 9 and 8.5 Ma. Between both depocentres there is an ENE-trending transfer zone (Radic et al. 2002). The volcanic rocks exposed in the western part of the Neuquen Embayment, such as those in Cerro Cabras, Cerro Tormenta, Desfiladero Negro and Cerro Sur de Los Overos (at approximately 37°30'S), are alkaline basalts that have

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ages between 23 and 19 Ma, Oligocene-early Miocene (Ramos & Barbieri 1989). These rocks have very low initial 87Sr/86Sr ratios of the order of 0.7035, smaller than the present magmatic arc (Ramos & Barbieri 1989) and are interpreted as having been formed in an extensional setting, with no influence from the subducted slab (Kay 2001, 2002). These low initial ratios of 87Sr/8 Sr associated with high neodymium isotopic ratio eNd (>+5), have been reported as far south as the 41°S latitude at both sides of the Principal Cordillera by Munoz et al. (2000). Several authors have emphasized the large volume of volcanic rocks erupted in a short time span during late Oligocene-early Miocene times (Jordan et al. 2001; Munoz et al. 2000; Folguera et al. 2003£). This fact, together with the unusually primitive nature of the magmas at these latitudes (Lopez Escobar & Vergara 1997), indicate an important extension of forearc, arc and intra-arc regions at that time. This extensional regime is linked with an important shift to the trench of the magmatic activity after the late Eocene. Middle-late Miocene arc The eruption of the Trapa-Trapa and Mitrauquen Formations along the axis of the Neuquen Andes between 18 and 8 Ma coincides with the emplacement of middle-late Miocene stocks of granodioritic composition with ages between 16 and 10 Ma (Moreno & Parada 1976; Munizaga et al. 1985). The emplacement of these stocks is coeval with a new pulse of expansion of the volcanic activity in the foreland. This broadening of the magmatic activity was recorded mainly to the north of the Cortaderas lineament in the Sierra de Huantraico (37°S), where hornblende andesites and dacites were erupted, like the Pichi Tril Andesite at about 18 ± 2 Ma (K-Ar whole rock: Ramos & Barbieri 1989). Near Cerro Bayo, on the eastern edge of the Huantraico syncline and at Filo Morado on its NW edge, Cobbold & Rossello (2003) have sampled lava flows, obtaining early Miocene ages of 22.1 ± 0.5 and 22.2 ± 0.2 Ma, respectively, by Ar/Ar on whole rock. There are several other centres with comparable characteristics in the Sierra de Huantraico and further north. These andesitic rocks extend north of Rio Colorado up to the Sierra de Chachahuen where they have been studied by Kay (2002). The hornblende-bearing andesite of this locality are about 480 km away from the trench. This important period of broadening of the magmatism was associated with several

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Fig. 6. Palaeogeographic map of the Cura Mallin Basin with indication of the early Miocene magmatic front, the extension of the outcrops and available ages of the interbedded volcanic rocks.

subvolcanic centres. The associated volumes of volcanic material were far less than those produce during the Oligocene-early Miocene period described above. The geochemistry of these rocks indicates a normal magmatic arc with hornblende-bearing calc-alkaline rocks. Several authors recognized an important period of deformation between 16.3 ±0.1 (Ar/ Ar) and 6.7 ± 0.5 Ma (K-Ar) by Kozlowski et al (1996) in the Coyuco syncline, north of Huantraico, where deposits of latest Miocene

age and Pliocene are not folded. These values are similar to the constraints proposed by Ramos & Barbieri (1989) for the folding of the lavas and pyroclastic rocks of the southern end of Huantraico bracketed between 18 and 9 Ma. The deformation in the inner sector of the Cura Mallin Basin was constrained between 8 and 5 Ma on the basis of the unconformity that separates folded products of the Cura Mallin Basin and Pliocene volcanic rocks (Folguera et al. 20036).

TECTONIC EVOLUTION OF THE NEUQUEN ANDES

Pliocene-Pleistocene arc and intra-arc basin A new period of intense volcanic activity and migration to the trench of the volcanic front started in the early Pliocene. The widespread magmatism of this period led Mufioz & Stern (1985, 1988) to recognize two belts of volcanic rocks. The volcanic front along the axis of the Principal Cordillera and an extensive belt with intra-arc volcanics, both of which are characterized by poorly evolved lavas with low initial 87 Sr/86Sr ratios (of the order of 0.70380.7040), erupted in an extensional regime. The Pliocene rocks were assigned to the Cola de Zorro Formation by Vergara & Mufioz (1982), and consist of volcanic breccias and lavas of basaltic-dacitic composition. This volcanic sequence records large volumes of erupted material and numerous evidence for synextensional deposition (Folguera et al 2003&). Rapid thickness variations from 1900 m to a few tens of metres, changes in the polarity of the half-graben system, as well as syndepositional discontinuities indicate generalized extension in this period. These volcanic rocks were erupted between 5 and 3 Ma. These rocks covered most of the Loncopue trough and extend along the axis of the Principal Cordillera, west of the drainage divide. Similar basalts in Paso Pichachen have been dated at 3.6 ± 0.2 and 3.6 ± 0.5 Ma (Munoz Bravo et al. 1989) and in Paso Pino Hachado at 4.8 ± 0.2 Ma (Linares & Gonzalez 1990). Coeval with this Pliocene eruption, there are isolated cones and lava flows of alkaline basalts developed in the foreland region along the northern end of the Sierra de Los Chihuidos that have been described by Ramos (1981). These volcanic rocks erupted in the western sector of the Neuquen Embayment, such as in the Cerro Parva Negra volcano, and were dated at 4.5 ± 0.5 Ma (Ramos & Barbieri 1989). Further east, in Aguada Rincon and in Cerro La Manea, alkaline basalts described by Holmberg (1964) in the foothills of the Auca Mahuida volcano yielded ages of 4.8 + 2 and 3.4 + 0.5 Ma at the base of the volcanic sequence (Valencio et al 1979). These alkaline rocks in the retro-arc have been attributed to a mild within-plate extension without relationship to the subducted slab (Kay 2001). A new period of volcanic activity took place between 2 and 0.5 Ma. The erupted volcanic rocks were concentrated in a narrower belt than those of the previous pulse (Folguera et al. 20036; Melnick et al. 2003, 2005). Rocks of this new event have been recognized in Laguna del Barco area, west of Copahue volcano, with

25

ages ranging from 2.68 to 2.60 Ma (Melnick et al. 2005). They are widespread in Rio Pino Solo and Piedra Blanca, east of Paso Pino Hachado, with several ages ranging from 1.40 ± 0.2 to 1.6 ± 0.2 Ma (K-Ar whole rock: Mufioz & Stern 1988). There is a spatial coincidence between the early Pliocene and late Pliocene-Pleistocene volcanic fronts, as denoted by Munoz Bravo et al. (1989) and Lara et al. (2001), but the volume of eruption and the area are more restricted in the younger event. During the Pliocene-Pleistocene important volcanic activity was registered in the retro-arc. The Auca Mahuida Volcano, located 500 km from the trench, erupted through a series of abundant but small monogenic centres extruding large amounts of basaltic lavas ranging in age from 1.7 + 0.2 to 0.9 ±0.07 Ma (Ar/Ar, plateau and isochron ages: Rossello et al. 2002). The last activity is interpreted to be related to the trench migration of the late PleistoceneHolocene volcanic front described by Mufioz & Stern (1988), in which a 30-50 km displacement is reported. At this time, a reactivation of the Loncopue trough controlled the eruption of many monogenic basaltic cones and small lava flows as seen west of Loncopue town and along the foothills of the Neuquen Andes.

Magmatic and tectonic styles The alternation of periods of voluminous arc magmatism and intra-arc basin development with intervals of reduced arc magmatism and deformation has captured the attention of several previous authors (e.g. Folguera et al. 2002). However, the mechanism and causes of such links are still poorly understood. Several hypotheses have been advanced, mainly to explain the voluminous magmatism associated with intra-arc development. Munoz & Stern (1988) proposed thermal or mechanical perturbations of the subcontinental mantle associated with subduction. These perturbations were interpreted as being the result of diapiric mantle upwelling or some other process of lithospheric thinning and erosion associated with continental extension. Other authors have proposed that extension could be a consequence of important strike-slip displacement of the Liquine-Ofqui Fault Zone (McDonough et al. 1997; Suarez & Emparan 1997). Munoz Bravo et al. (1997) considered that melting during Oligocene-early Miocene was caused by asthenospheric upwelling driving the crustal extension, rather than slab dehydration in the asthenospheric wedge. The presence of alkali basalts along the Central Valley,

26

V. A. RAMOS & A. FOLGUERA

comparable with those erupted in intra-arc and retro-arc basins associated with the active arc products, suggests derivation from an oceanictype mantle unmodified by components derived from a subducted slab. The asthenospheric up welling during the late Oligoceneearly Miocene times was explained by Munoz et al (2000) and Stern et al (2000) as a consequence of an asthenospheric window developed during a period of plate reorganization. Based on the widespread seismic evidence for extension Jordan et al. (2001) proposed that, instead of localized transtension, the bulk strain was horizontal extension. The abnormal melting in the Central Valley 170km from the trench may reflect an increased flux of water as a consequence of more rapid subduction and this high pore pressure may have favoured a decrease in forearc topography. The same authors also recognized that increased heat flux would produce uplift and, in turn, would provoke moderate extension. This abnormal heat flux could be related to a transient hot spot as proposed further east at 27 Ma by Kay et al. (1993). The link between rapid subduction and extension during the Oligocene-early Miocene, followed by later shortening, was challenged by Stern et al. (2000) and Godoy (2002). This last author presented evidence that along the Chilean margin there were segments that did not record any extension at that time, as well as others with extension and no subsequent shortening. There are several facts that should be considered when trying to understand the alternation of different magmatic and tectonic styles along the Andes at these latitudes. First, that this alternation is almost unique along the Central Andes, and therefore the explanation should have some exceptional causes. Secondly, that although extension was generalized at these latitudes in the forearc (Cisternas & Frutos 1994), arc and intra-arc (Jordan et al. 2001; Radic et al. 2002; Folguera et al. 2003<2, b), and a mild extension in the retro-arc (Kay 2001), subsequent shortening was not that important (Ramos et al. 2004). The common magmatic and structural features of the alternating Mesozoic and Cenozoic settings can be summarized in the following twostage process. Intense intra-arc magmatism and extension Large volumes of igneous activity and poorly evolved magmas, located at a short distance (typically less than 170km) from the trench, are recurrent in the Jurassic-Early Cretaceous, Oligocene-early Miocene and Pliocene times. The presence in extreme cases, as in the

Oligocene, of alkali basalts, typical of retro-arc settings together with arc products, were interpreted as the result of an input of hot and undepleted asthenosphere, different from a typical asthenospheric wedge. Geochemical characteristics show flat RREE patterns with low La-Yb ratios, which is common for low-pressure crystallization, typical of a thinned crust (Munoz & Stern 1988). These magmatic episodes were produced during periods in which the magmatic activity retreated toward the trench. The periods are associated with large areas of eruption that are significantly different to the normal, narrower volcanic arcs. At this time the volcanic front was associated with wide intra-arc basins and, in certain cases, with retro-arc alkali magmatism. The increase in alkali contents, mainly potassium (Kay et al. 2005), is interpreted as evidence of decreased melting toward the foreland. Where subsurface data are available, or when the palaeogeography can be reconstructed, well-defined half-graben systems are recognized as controlling the eruption of large volumes of lava associated with alluvial and lacustrine deposits. Evidence of syntectonic deposition has been observed in lavas and interfingered sediments as growth discontinuities and rapid changes in thickness. The structural trend of the half-grabens resulted from pre-existing basement fabrics or the changing orientation of the stress field as a consequence of convergence vector adjustments (Fig. 7). Extension at those times was generalized, affecting a great part of the continental margin from the forearc to the foreland. The greatest stretching was concentrated along the axis of the cordillera with mild extension detected in the foreland. Magmatic arc shifting and deformation The periods of major volcanic activity described above alternated with periods in which narrow belts of plutono-volcanic complexes were formed. The main difference is the presence of coarse-grained tonalites and granodiorites that implies important uplift and consequent denudation after crystallization. The volcanic rocks of these periods are more evolved products that include dacites and andesites as the dominant rock types and indicate important differentiation from parental magmas, either by fractional crystallization or by assimilation of the crust. Geochemical characteristics show the typical trends of calc-alkaline rocks, with steep RREE patterns and higher La-Yb ratios, formed at a higher pressure, probably in a thickened crust.

TECTONIC EVOLUTION OF THE NEUQUEN ANDES

27

Fig. 7. Conceptual cross-section illustrating the processes related to the steepening of the subduction zone and related magmatism.

The magmatic belts indicate a conspicuous foreland migration with respect to the previous arcs, as observed during the Late Cretaceous Palaeogene, and from middle Miocene to late Miocene times (Fig. 8). The volcanic products were emplaced in previously deformed rocks. For example, the Late CretaceousPalaeogene intrusives have deformed Cretaceous deposits as country rocks. Overall the magmatic deposits are associated with synorogenic deposits such as the clastic sequences of red beds and conglomerates of the Neuquen Group (Late Cretaceous), and the elastics and carbonates of the Malargiie Group (MaestrichtianPaleocene).

The magmatic front migrated towards the foreland more than 350 km away from the trench, and there was no magmatic activity along the previous magmatic axis, or it is punctuated by the emplacements of plutonic stocks of coeval granitoids. The lack of significant subduction erosion at these latitudes, as postulated by Ramos (1988), Stern (1991) and Kay et al (2005), precludes crustal erosion in the forearc as a mechanism of arc migration toward the foreland. The decrease in magmatic volume, together with the expansion of the magmatic activity to the foreland, can be better explained by changes in the Benioff geometry.

Fig. 8. Conceptual cross-section illustrating the processes related to the flattening of the subduction zone and related magmatism.

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Integrated model In order to explain the alternation of periods with intense intra-arc and arc magmatism and volcanic front retreats, with periods of arc magmatism and foreland expansion and deformation, an integrated model based on the processes presented by James & Sacks (1999) will be analysed. These authors emphasized the profound tectonic and magmatic effects that result from the interaction of hot asthenosphere and hydrated mantle during the transition from flat to normal subduction. Steepening of the angle of subduction brings an influx of hot asthenospheric mantle from depth to fill the opening mantle wedge. When this happens in a thin crust, such as the present setting of the Neuquen Basin, the process will induce large amounts of melting in the asthenospheric wedge and crustal extension with the eruption of large volumes of poorly evolved magmas. In contrast, during a shallowing of the subduction zone, the mantle wedge closes progressively during flattening and asthenospheric material is expelled eastward. Volcanism is shut off along the main volcanic arc (Kay et al 1991). Such a process may account for the observed distribution of magmatic rocks and extension in

the Neuquen Andes since the Jurassic (Fig. 9). Variations in the angle of subduction in the different Benioff zones can explain both the observations from the main arc and also the interactions observed between uplift and collapse of the Principal Cordillera and the tectonic effects in the Neuquen Embayment. The variations in the angle of subduction are generally attributed to changes in slab buoyancy, probably by subduction of thickened oceanic crust and retardation of the basalt-eclogite transition in the young and relatively hot oceanic plate (Pilger 1984; James & Sacks 1999; Gutscher 2002). The different cycles of intense magmatism and extension, alternating with foreland shifting of the magmatism and compression, have general similarities in the Neuquen Andes in the last 200 Ma. However, a detailed analysis shows some striking differences generated by the changing kinematics of the interaction between the different Pacific plates and the continental plate. As a result of these changes each cycle has some important peculiarities superimposed on the general trend of steepening and shallowing of the subduction zone.

Fig. 9. View to the east of the Late Cretaceous angular unconformity between the upper member of the Rayoso Formation (Kr) and the red beds of the Candeleros Formation (Kc). The Candeleros Formation is the lower unit of the Neuquen Group east of Cerro Rayoso on the western margin of Rio Neuquen.

TECTONIC EVOLUTION OF THE NEUQUEN ANDES

Tectonic interaction of the Andes and the Neuquen Basin The alternation of periods of intense deformation with others of extension has received considerable attention since plate tectonic concepts were first applied to the formation of orogenic belts (Charrier 1973). However, these changes can be addressed only through the recent understanding of the sublithospheric processes, based on geophysical, petrological and tectonic studies. On these bases, the interaction between the processes observed in the main cordillera and the effects in the Neuquen Basin will be examined. Jurassic-Early Cretaceous stage Although rocks of this stage are poorly exposed in the Neuquen Andes at these latitudes, there is enough information at both ends of the segment. The magmatic activity occurred in an extensional regime (Munoz 1984; De la Cruz & Suarez 1997). Owing to the onset of subduction in the earliest Jurassic (Kay 1993; Franzese & Spalletti 2001), the extensional tectonic regime was characterized by pulses of negative rollback velocity in the trench. As a consequence of this, the Neuquen Basin was affected by generalized rifting enhanced along pre-existing crustal weakness zones, such as the Huincul Fault Zone. Several pulses of horizontal extension, partitioned by the orientation of the weakness zones, resulted in pure extension and localized transtension-transpression. Detailed examples of this interaction have been presented from along the Huincul Fault Zone by Vergani et al (1995), Pangaro et al (2002) and Mosquera (2002). This widespread rifting was also depicted in several parts of the basin by Zapata et al. (2002). Late Cretaceous-Palaeogene stage The emplacement of the Late Cretaceous batholith was followed by the migration of the magmatic arc to the foreland. This migration is associated with the deformation of the main Andes as it is observed in the 6 km uplift of Cordillera del Viento, which exposed coarsegrained granitoids prior to the Paleocene volcanics. Deformation is also evidenced by growth strata in the red beds of the Diamante Formation. These facts clearly indicate a period of progressive deformation that started in the Late Cretaceous and continued until the late Eocene. The beginning of compressive deformation was produced by tectonic inversion of pre-existing

29

normal faults during a period of shallowing of the Benioff zone. The Agrio Fold and Thrust Belt was formed at this time, with a combination of inversion tectonics in the inner sectors (Vergani et al. 1995; Ramos 1998) and localized thin-skinned tectonics in the outer areas (see Zapata & Folguera 2005). As a result of this, the first foreland basin was formed at these latitudes and filled with the synorogenic deposits of the Neuquen Group. The angular unconformity with the pre-orogenic deposits is seen between the top of Rayoso Formation and the base of Neuquen Group along the Chihuidos Ridge east of Rio Neuquen (Fig. 9). Compelling evidence for a Late Cretaceous deformation was also presented by Cobbold & Rossello (2003). This deformation was associated with a high rate of orthogonal convergence in the Late Cretaceous (Larson 1991) that became quite oblique during the Paleocene (Pardo Casas & Molnar 1987). Cobbold et al. (1999) interpreted the structures formed during the Eocene deformation as mainly transpressive. New fissiontrack data (Grafe et al. 2002) indicate that north of 39°S, the axial part of the main cordillera was uplifted during the Eocene. In contrast, immediately to the south, the northern extreme of the intra-arc Liquine Ofqui Fault Zone (Fig. 3) has imprinted younger deformations on the highest parts of the cordillera. Palaeogene synorogenic deposits are more developed in southern Mendoza than in Neuquen where they are restricted north of the Cortaderas lineament (Ramos 1981). However, the Malargiie Group has a depositional system controlled by the flexural subsidence as a result of tectonic loading of the Principal Cordillera (Tunik 2000). Oligocene-early Miocene stage The retreat of the magmatic activity toward the trench, associated with the steepening of the subduction zone, produced the Cura Mallin Basin and the generalized extension of the Pacific continental margin as described by Cisternas & Frutos (1994). This extensional process has been described from the latitudes of the present study down to the 41°S latitude (Munoz et al. 2000). The injection of hot asthenosphere from the foreland sublithospheric mantle to the newly formed asthenospheric wedge produced a large amount of melting and the formation of arc products and alkali basalts near the Central Valley closer to the trench than in previous periods. Large volumes of poorly evolved magmas produced the Oligocene-early Miocene volcanics in the Neuquen Andes.

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At that time, the Loncopue trough was the down-thrown block of a normal fault with almost 3 km of displacement located west of Andacollo (Jordan et al. 2001). Except for the generation of a series of extensional basins, such as the Cura Mallin, Collon Cura and Nirehuao, along the foothills of the Neuquen Andes, as described by Gonzalez Diaz & Nullo (1980) and Dalla Salda & Franzese (1987), most of the Neuquen Basin only records a mild subsidence with deposition of fall tuffs of Palaeogene age. Middle-late Miocene stage This was the second period of deformation of the Neuquen Basin. The present structures of the Agrio Fold and Thrust Belt were formed at this time, as well as the final uplift of the Chihuidos ridge, and the Anelo synorogenic depocentre was formed east of the Chihuidos. The deformation was dated at less than 16 Ma by Kozlowski et al. (1996) north of Huantraico. Localized synorogenic deposits, such as the Tralalhue Conglomerates (Ramos 1998) west of Rio Neuquen and other depocentres, show the limited subsidence of the western part of the embayment during this period of compressional deformation. This period coincides with an important flattening of the subduction that recorded arcrelated volcanic rocks of late Miocene age in the Sierra de Chachahuen (Kay 2002), almost 500 km away from the trench. The Neuquen Andes at this time recorded an important deformation of the Cura Mallm Basin (Folguera et al. 2003/7, 2004) controlled by tectonic inversion of previous normal faults. Although total contraction is not as important as in the northern segments of the Central Andes (Ramos et al. 2004), the maximum shortening of the Neuquen Andes was attained at this stage. Pliocene-Pleistocene stage This period is characterized by steepening of the flat subduction proposed by Kay (2002) in the late Miocene in the Chachahuen region. As the flat subduction was at a maximum north of Cortaderas lineament, this area recorded the maximum extension and magmatic activity as denoted by Munoz Bravo et al. (1989). The northern areas have not only the arc and intraarc basin developed, but also important retroarc volcanoes such as the Tromen and Auca Mahuida. The retreat of the magmatic front controlled the development of the half-graben system filled by the volcanic rocks of the Cola de

Zorro Formation. The early Pliocene corresponds with the maximum horizontal extension and a large volume of volcanic rocks were extruded (Folguera et al. 2004). From the early to late Pliocene-Pleistocene the area of maximum subsidence was shifted to the east, to the present Loncopue trough. This time marks the maximum development of the Loncopue trough, and neotectonic features indicate localized transtension such as the collapse of Caldera del Agrio and other rhomboidal pull-apart basins observed in the foothills (Folguera & Ramos 2000; Melnick et al 2005). Contractional deformation related to the Liquine-Ofqui Fault Zone during the Pleistocene deformed the northern segment of the Loncopue trough. The Plio-Pleistocene volcanic arc is currently being cannibalized by the thrust front north of 37°30/ S, as indicated by recent neotectonics (Folguera et al. 2004). Along the axis and on the western slopes of the cordillera there is evidence of extension (e.g. in Laguna La Laja), where seismic sections show negative flower structures, along transtensional segments (Melnick et al. 2003). The southern sector of the Loncopue trough, south and east of the Loncopue town, shows evidence of extensional neotec tonics along the eastern margin of the trough, although there are no available subsurface data. Conclusion The analysis of the expansions and retreats of the magmatic arc through time shows a strong control over the tectonic regime of the Neuquen Embayment. As a result of this, two periods of progressive deformation to the foreland can be identified. The first period of contraction post-dated the generalized extension that lasted from the Early Jurassic to the Early Cretaceous. A progressive deformation started with the emplacement of the Late Cretaceous granitoids in the arc and the foreland migration of the magmatism that lasted until the late Eocene. Late Cretaceous deformation began with orthogonal contraction and ended in the Palaeogene with generalized transpression (Cobbold & Rossello 2002). The second contractional deformation began in the middle Miocene and lasted until the late Miocene. Both periods of contraction were linked to shallowing of the subduction zone as postulated by Kay (2002) for the late Miocene (Fig. 10). The two compressional periods were followed by steepening of the subduction zone, in the Oligocene and the Pliocene. This resulted in hot asthenosphere from the subcontinental mantle being injected into the asthenospheric wedge,

TECTONIC EVOLUTION OF THE NEUQUEN ANDES

31

Fig. 10. Summary of changing geometries of the palaeo-Benioff zones through time, based on the magmatic evidence and structural styles. See discussion in the text.

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Neuquen basin, Argentina. Tectonophysics, 314, 423-442. COUCH, R., WHITSETT, R., HUEHN, B. & BRICENOGUARUPE, L. 1981. Structures of the Continental Margin of Peru and Chile. Geological Society of America, Memoirs, 154, 703-726. DALLA SALDA, L. & FRANZESE, J. 1987. Las megaestructuras del Macizo y la Cordillera Norpatagonica argentina y la genesis de las cuencas volcanosedimentarias terciarias. Revista Geologica de Chile, 31, 3-13. DALY, M. 1989. Correlations between Nazca/Farallon plate kinematics and forearc evolution in Ecuador. Tectonics, 8, 769-790. DECELLES, P. & GILEST, K. 1996. Foreland basin systems. Basin Research, 8, 105-123. DE LA CRUZ, R. & SUAREZ, M. 1997. El Jurasico de la cuenca de Neuquen en Lonquimay, Chile: Formation Nacientes del Biobio (38°-39°S). Revista Geologica de Chile, 24, 3-24. DOMINGUEZ, E.A., ALIOTTA, G., GARRIDO, M., DANIELI, J.C., RONCONI, N., CASE, A.M. & PALACIOS, M. 1984. Los Maitenes-El Salvaje. The field work of this study was made through funding Un sistema hidrotermal tipo porffrico. IX° Confrom PICT 06729/99 of the Agencia Nacional de Promogreso Geologico Argentino (Bariloche), Actas, 7, cion Cientifica y Tecnologica for the 'Evolution tectonica 443-458. y paleogeografica de los Andes Centrales'. The critical FOLGUERA, A. & RAMOS, V.A. 2000. Control estruccomments and reviews of T. Sempere and P. Cobbold have contributed to a better comprehension of the manutural del Volcan Copahue: implicancias tectonicas script. The authors acknowledge the members of the para el arco volcanico cuaternario (36°-39°S). Laboratorio de Tectonica Andina of the UBA for logistic Revista de la Asociacion Geologica Argentina, support and critical review. 55, 229-244. FOLGUERA, A., RAMOS, V.A. & MELNICK, D. 2002. Partition de la deformation en la zona del arco volcanico de los Andes neuquinos (36°-39°S) en los ultimos 30 millones de anos. Revista Geologica References de Chile, 29, 227-240. ALLMENDINGER, R.W., JORDAN, T.E., KAY, S.M. & FOLGUERA, A., ARAUJO, M., RAMOS, V.A., MELNICK, ISACKS, B.L. 1997. The evolution of the AltiD., HERMANNS, R., GARCIA MORABITO, E. & plano-Puna Plateau of the Central Andes. Annual BOHM, M. 2003a. Seismicity and variation of the Reviews of Earth Planetary Sciences, 25, 139-174. crustal tensional state of the retro-arc in the BOHM, M., LOTH, S., ECHTLER, H., ASCH, G., southern central Andes during the last 5 Ma BATAILLE, K., BRUHN, C, RIETBROCK, A. & (37°30/-39°S). X° Congreso Geologico Chileno WIGGER, P. 2002. The Southern Andes between (Conception), Actas, electronic files. 36° and 40°S latitude: seismicity and average seismic velocities. Tectonophysics, 356, 275-289. FOLGUERA, A., RAMOS, V.A. & MELNICK, D. 2003£. Recurrencia en el desarrollo de cuencas de CHARRIER, R. 1973. Interruptions of spreading and the intraarco, Cordillera Neuquina (37°30/). Revista compressive tectonic phases of Meridional Andes. de la Asociacion Geologica Argentina, 58, 3-19. Earth and Planetary Science Letters, 20, 212-249. CISTERN AS, M.E. & FRUTOS, J. 1994. Evolucion FOLGUERA, A., RAMOS, V.A., HERMANNS, R.L. & NARANJO, J. 2004. Neotectonics in the foothills tectono-estratigrafica de la Cuenca terciaria de los of the southernmost central Andes (37°-38°S): Andes del Sur de Chile (37°30/-40°30/ Lat.S). Evidence of strike-slip displacement along the 7° Congreso Geologico Chileno (Conception), Antinir- Copahue fault zone. Tectonics, 23, Actas, 1, 6-12. TC5008, doi: 10.1029/2003TC001533. COBBOLD, P.R. & ROSSELLO, E.A. 2002. Phases of Andean deformation, foothills of Neuquen basin. FRANCHINI, M.B., LOPEZ ESCOBAR, L., SHALAMUK, I.B.A. & MEINERT, L.D. 2003. Paleocene, calcIn: 5th International Symposium on Andean Geodyalkaline subvolcanic rocks from Nevazon Hill namics, Toulouse. Extended Abstracts. IRD, Paris, area (NW Chos Malal Fold Belt), Neuquen, 153-156. Argentina, and comparison with granitoids of the COBBOLD, P.R. & ROSSELLO, E.A. 2003. Aptian to Neuquen-Mendoza volcanic province. Journal of recent compressional deformation of the Neuquen South American Earth Sciences, 16, 399-422. Basin, Argentina. Marine and Petroleum FRANZESE, J.R. & SPALLETTI, L.A. 2001. Late Triassic Geology, 20, 429-443. continental extension in southwestern Gondwana: COBBOLD, P.R., DIRAISON, M. & ROSSELLO, E.A. 1999. Bitumen veins and Eocene transpression, tectonic segmentation and pre-break-up rifting.

inducing large amounts of melting and the formation of voluminous magmatism. Generalized extension prevailed at Oligocene times associated with oceanic plate reorganization and the break-up of the Farallones Plate. Extension in the Pliocene, although possibly related to a minor reorganization of the oceanic plates at about 5 Ma, progressively changed to transtension in late Pliocene and Pleistocene times, as proposed by Folguera et al (2003/?). These episodes of extension produced the Loncopue trough that attained its present morphology during the late Pliocene and Pleistocene. Active shortening is concentrated in the foothills north of 37°30'S. The alternation of short intervals of compression with minor shortening and periods of generalized extension produced the unique morphology of the Neuquen Andes.

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Tectonic evolution of the Andean Fold and Thrust Belt of the southern Neuquen Basin, Argentina T. ZAPATA1'2 & A. FOLGUERA2 l Repsol-YPF, Finding Team Faja Plegada, Esmeralda 255, of 1002, Buenos Aires C1035ABE, Argentina (e-mail: [email protected]) 2Laboratorio de Tectonica Andina, Departamento de Geologia, Universidad de Buenos Aires 1428 Buenos Aires, Argentina (e-mail: [email protected]) Abstract: The Andean Fold and Thrust belt between 36° and 39°S can be divided in two sectors. The Eastern Sector corresponds to the Agrio Fold and Thrust Belt (FTB) characterized by a major exhumation during the Late Cretaceous, and minor deformation during the late Eocene and Late Miocene. The Western Sector corresponds to the main cordillera and is characterized by a complex evolution that involves periods of out-of-sequence thrusting with respect to the previously deformed outer sector, and pulses of relaxation of the compressive structure. Cretaceous uplift constituted an orogenic wedge that extended to the inner sectors of the Agrio FTB. Eocene compression was mainly concentrated within the Western Sector but may have reactivated the pre-existing structures of the Agrio FTB, such as the Cordillera del Viento. Late Miocene minor compressional deformation occurred in the retro-arc area and extended into the foreland area. This deformation event produced the closure of a short-lived intra-arc basin (Cura Mallin Basin, 25-15 Ma) at the innermost sector of the FTB. The Pliocene and Quaternary, between 37°30'and 39°S, have been periods of relaxation of the inner part of the FTB and fossilization of the Agrio Fold and Thrust Belt. Localization of episodic late Oligocene-Early Miocene and Pliocene to the present extensional structures in the intra- and inner retro-arc is controlled by pre-existing Jurassic half-grabens related to the formation of the Neuquen Basin. The Jurassic rift seems to be controlled by deep crustal-lithospheric discontinuities derived from a ProterozoicPalaeozoic history of amalgamation in the area, now deeply buried under multiple episodes of Mesozoic-Tertiary synorogenic and synextensional sedimentation.

Geophysical studies have revealed that the Andes mountain belt is extremely variable in crustal thickness and topography (Introcaso et al 2000). The topography varies between broad amplitudes greater than 700 km measured from the trench, and narrow belts restricted to the inner sectors of the fold and thrust belt. These variations are mainly related to shortening within the Andes (Ramos et al. 2004). However, the causes of variable shortening and relief remain open to discussion and include several key factors: (1) shortening of the mantle lithosphere related to overthrusting of the Andes over old cratonic shields (Lyon-Caen et al. 1985; Lamb & Hoke 1997; Kley et al. 1999); (2) pre-existing anisotropies in the foreland of the orogenpre-dating the Andean orogeny, which differentially deformed under compression (Allmendinger & Gubbels 1996); (3) changes in the lithospheric thermal structure and the consequent development of brittle-ductile transitions

that become new detachments where the upper crust yields and is stacked over the foreland (James & Sacks 1999; Ramos et al 2002); and (4) climate (Beaumont et al. 1992; Thomson 2002). At this latitude (37°-39°S), the Andean mountain belt deforms the Mesozoic Neuquen Basin (Fig. 1). The maximum topographic heights are restricted to a narrow band next to the volcanic arc, and the external zone represents a smooth surface where older deformations have taken place during the Late Cretaceous and Palaeogene (Fig. 1) (Zapata et al. 1999, 2002). The lack of amplitude and height of the orogenic system, in comparison with neighbour segments to the north, are in accordance with the minimum shortening computed from surface structures and crustal roots (Zapata et al. 1999; Ramos et al. 2004). The tectonic evolution of the Andean mountains in the southern portion of the Neuquen Basin reveals certain anomalies to the general

From\ VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Cas Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 37-56. 0305-8719/05/$15.00 © The Geological Society of London 2005.

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Fig. 1. Regional location map, where main rnorphostructural units of the Andes between 36° and 40°S are displayed. (A) Arc and retro-arc rnorphostructural units. (B) Fore arc to retro-arc systems mentioned throughout the paper. (C) Thematic mapper scan of the area occupied by the Neuquen Embayment during the Mesozoic from the western to the eastern side of the present Andean belt. The square represents the area of Figure 5 and the black line indicates the position of the profile in Figure 10.

picture of progressive foreland-propagatingatinh deformation in the Andes. These anomalies are characterized by a positive roll-back velocity, since the break-up of southern Gondwana (Ramos 19996), and periods of foreland propagation of thrust sheets alternating with periods of tectonic relaxation, possibly originated from

changes in the Wadati-Benioff geometry. However, are these real anomalies from the Andean orogeny point of view or do they exemplify a long-standing process in many segments along the Andean chain that elsewhere have been obscured by younger tectonic imprints? Other subduction-related orogens around the

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world show that tectonic relaxation in the inner sectors of fold and thrust belts is a common process, and is related to changes in boundary conditions such as oblique convergence between plates, subduction acceleration and roll-back velocity (Pe-Piper et al 1995; Petford & Atherton 1995; Jolivet & Faccenna 2000; Morley 2001; Melnick et al 2003). The objective of this study is to illustrate the evolution of a particular segment of the Andes, where the presence of certain anomalous features may highlight the evolutionary path of the whole Central Andean system. This segment (37°30/39°S) is suitable to reveal these anomalous features, particularly along the inner sectors of the fold and thrust belt, due to the oscillatory behaviour of the arc front (Mpodozis & Ramos 1989), in contrast to the rest of the Austral Central Andes (5°-35°S) where the continuous foreland progression of the volcanic arc has obscured parts of the contractional history. This study focuses on the connection of the foreland and hinterland deformation, and assesses the complexity of the process by addressing the questions posed above. Tectonic framework of the Andean Fold and Thrust Belt Three morpho-structural units can be identified between 36°I5 f and 38°30/S (Figs 1 & 2) from west to east. (1) The main Andean cordillera

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(11°45' - 71°W), located at the arc front and at the inner retro-arc area, where mainly Tertiary volcaniclastic sequences are exposed. These ranges, located mostly within the hinterland orogenic region, have experienced a rather complicated evolution where the normal progression of thrusts sheets has been interrupted during periods of orogenic relaxation (Ramos & Folguera 2005). (2) The inner sector of the Agrio Fold and Thrust Belt (FTB) is composed of Jurassic-Lower Cretaceous predominantly marine units of the Neuquen Basin that have been deformed since Late Cretaceous times. (3) The outer sector of the Agrio FTB, where mainly Cretaceous-Miocene rocks occupy an almost peneplaned surface, that was continuously deformed up to the late Miocene (Figs 1 & 2). The Agrio FTB The Neuquen Basin Mesozoic deposits are incorporated in both the Inner and the Outer Agrio FTB (Ramos 1977). Documentation of timing of tectonic activity by synorogenic Tertiary volcaniclastic deposits is restricted to a few remnants in narrow intermontane basins (Ramos 1998; Zapata et al. 2002) (Figs 3 & 4). Stratigraphy. The basement of the Mesozoic Neuquen Basin is composed of a thick pile of volcanic sequences

Fig. 2. Lithospheric cross-section representing the Agrio FTB and inner areas of d seismicity. The forearc seismicity was taken from Bohm et al (2002) and the retro-arc seismicity from Folguera et al. 2003.

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of the Choiyoi Group (Fig. 3). These units are related to an episode of crustal extension that affected most of the Andean region and Patagonian platform during the Permian (Kay et al. 1989). These rocks crop out at the NE and SW basin border, where the Late Jurassic-Early Cretaceous tectonic subsidence that formed the Neuquen Basin did not take place; and along the inner sector of the Agrio FTB (Cordillera del Viento, Fig. 1), (Zollner & Amos 1973) as part of uplifted basement blocks during the Andean deformation (Kozlowski et al. 1996; Ramos 1998; Zapata et al 1999, 2002).

The oldest Mesozoic units exposed on the western side of the fold and thrust belt correspond to the Kimmeridgian continental sandstone of the Tordillo Formation (Gulisano & Gutierrez Pleimling 1994). The older Jurassic sequence and basement units crop out at the core of the Cordillera del Viento range (Fig. 1), and has been documented by borehole data from the Cerro Mocho x-1, drilled by YPF (the former National Oil Company). Coeval Jurassic volcanic units west of the Cordillera del Viento are dated at 167.7 ± 8 Ma (Rovere 1993, 1998).

Fig. 3. Stratigraphic chart corresponding to the western slope of the Andes between 36° and 39°S.

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Fig. 4. (A) Agrio and Chos Malal FTB to the right and the main Cordillera to the left and main episodes of deformation registered by fission-track data, fossil associations in synorogenic strata and regional unconformities (ages were taken from Suarez & Emparan 1997; Ramos 1998; Zapata et al. 1999, 2002; Jordan et al 2001; Grafe et al. 2002). (B) Shaded topographic model. Plan view of the area represented in (A), where main tectonic elements are displayed with their corresponding morphological expression.

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The Jurassic sandstones are overlain by Tithonian-Neocomian marine shales, platform carbonates, salt and gypsum units (Huitrin Formation) of the Mendoza Group (Fig. 3). These successions represent the period of basin expansion throughout the whole retro-arc region. At this stage, the Neuquen Basin was a back-arc basin strongly affected by tectonic and eustatic sea-level changes due to a narrow connection with the ocean toward the NW (Ramos 1977). These Mesozoic units are capped by evaporites and continental sandstones - Huitrin Formation (Aguirre-Urreta & Rawson 1997). The Jurassic Auquilco and Lower Cretaceous Huitrin evaporites constitute the regional decollements (treated here as compressional detachment zones) (Fig. 3) of both the inner and outer Agrio FTB (Ploszkiewicz 1987). The westernmost outcrops, related to the Mesozoic extensional sequences of the Neuquen Basin, can be traced to the Quaternary arc on the Chilean slope of the Andean cordillera, where Lower Jurassic turbidites are thrust over Upper Miocene volcaniclastic sequences in the Lonquimay region (De la Cruz & Suarez 1997). The Mesozoic rocks were intruded along the eastern slope of the Andes by andesitic-dacitic subvolcanic bodies of the Collipilli Formation, which are unconformably overlaid by daciticandesitic lavas of the Cayanta Formation. These units are also known as Serie Andesitica (44_49Ma: Llambias & Rapela 1987, 1989). However, ages as old as Paleocene have been determined (Rovere 1998; Jordan et al 2001). In the Argentinean side, these volcanic facies regionally overlie the Mesozoic deposits with an angular unconformity (Fig. 3). These volcanics and related rocks are also found in a N-S trend along the western Neuquen Basin, although the western edge is not precisely defined due to the presence of profuse upper Palaeogene deposits. Structure The structure of the Agrio FTB is characterized by a combination of thin-skinned and thickskinned structures (Fig. 5). This FTB is bounded to the east by the Los Chihuidos high and to the west by the Loncopue trough (Fig. 1). The Agrio FTB is divided into two regions: the Inner Sector, where the exposed folds are related to the inversion of a Mesozoic extensional structure called the 'Tres Chorros extensional system', and the Outer Sector (Ramos 1977; Ramos & Barbieri 1989; Vergani et al. 1995), which is composed of tight axially extended anticlines that bound basement blocks (Fig. 5) (Zapata et al. 2002). The deformation in this part of the

fold and thrust belt has experienced several episodes (Zapata et al. 2002), that have been recorded by synorogenic deposits. The last compressional pulse of deformation is as old as Middle Miocene, when the whole FTB was thrust towards the foreland area, probably using the pre-existing Jurassic detachment (Zapata et al. 2002) inherited from the extensional period of the Neuquen Basin (Figs 1 & 2). The related synorogenic deposits of the Agrio FTB are buried in the Bajo de Anelo area to the east of the study area (Fig. 1) (Ramos 19990). Structure of the Inner Sector The Inner Sector of the Agrio FTB includes the southern extent of the 'Cordillera del Viento' basement uplift (Figs 1 & 4). It corresponds with the 'Tres Chorros extensional system' (Vergani et al. 1995) that is composed of a series of NW-trending broad anticlines associated with an inverted half-graben, inherited from the Jurassic Neuquen Basin opening (Ramos 1998) (Fig. 5). One such structure is the Cerro Mocho anticline (Fig. 5); a doubly-vergent basement uplift that resembles a 'pop-up' structure. Twodimensional (2D) seismic data show that the eastern limb of the structure is affected by a deep fault that cuts across the sedimentary sequence through the Auquilco Jurassic evaporites; transferring more than 6 km of shortening to the thin-skinned structures of the Outer Sector (Zapata etal 2002) (Fig. 5). Borehole data of the CMO x-1 well documented more than 1500 m of Jurassic synrift sequences (Fig. 5). This anomalous thickness is interpreted to be associated with an extensional half-graben that was inverted during the Andean orogeny. Structure of the Outer Sector The structure of the Outer Sector is composed of thin-skinned tight folds associated with deep faults (Fig. 5). Borehole and 2D seismic data show that the deep faults are detached from the Jurassic Auquilco evaporites and propagate up through the sequence until they reach the Cretaceous Huitrin evaporites (Fig. 5), conforming fault-bend fold structures. The upper units of these structures, corresponding to the Agrio Formation (Figs 3 & 5), have been locally deformed by flexural folding, adding a detachment folding component (Zapata et al. 2002). Borehole data show that the external structure of the Agrio FTB is characterized by a refolded triangle zone bounded on the eastern side by a backthrust that cores fault-related fold (Zapata et al. 2002) (Fig. 5).

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Fig. 5. Agrio FTB and the relationship between the inner and outer sectors. (A) Thematic mapper image and main localities cited in the text. (B) Integrated seismic line. (C) Structural section from the inner to the outer sector of the fold and thrust belt. Bas, basement of the Jurassic-Cretaceous succession Choiyoi - Group; Jc, Cuyo Group; Jt, Tordillo Formation, Ka, Agrio Formation, GN: Neuquen Group. CMOX-1: Cerro Mocho anticline. Pi: Pichaihue syncline. The black colour of the structural cross-section indicates the Huitrin evaporite (used as the upper detachment of the Agrio FTB structures).

Uplift and temporal constraints The beginning of the Andean deformation of the Inner Sector of the Agrio FTB took place during the Late Cretaceous (Ramos 1977). An Ar/Ar biotite cooling age of 69.09 ± 0.13 Ma from

igneous rocks of the Cordillera del Viento (northern continuation of the Inner Sector of the Agrio FTB, Fig. 4) (Kay 2001) supports the Late Cretaceous timing for uplift. The study also revealed that at that time there was already 3000 m of exhumation. Fission-track cooling

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ages from an igneous stock located at the core of the Cordillera del Viento and from sedimentary units from the eastern side recorded two main cooling events (Burns 2002): the first episode took place between 80 and 70 Ma (Late Cretaceous), and the second one between 7 and 5 Ma (Late Miocene). The Late Cretaceous cooling event coincides with a 71.5 ± 5 Ma K/Ar emplacement age for andesitic dykes of the Pelan Formation that intrude previously deformed Cretaceous sediments south of the Domuyo Volcano (Llambias et al 1978), 50 km to the north of the Cordillera del Viento (Fig. 4). An uppermost constraint for the Late Cretaceous deformation event of the Andean deformation is present on the western flank of the Cordillera del Viento. There an angular unconformity separates the Serie Andesitica igneous rocks from the Permian basement of the Choiyoi Group (Zollner & Amos 1973). Ar/Ar radiometric ages of the Serie Andesitica rocks yielded an age of 56 Ma (Jordan et al. 2001). The uplift of the Cordillera del Viento is interpreted to be related to a deep fault (bounded by a backthrust on its western side) that has transferred the shortening to the Outer Sector of the Agrio FTB (Kozlowski et al 1996). This fault may be connected to a previous Jurassic extensional detachment (reactivated as a decollement, i.e. compressional detachment, during the Andean deformation) that formed the Neuquen Basin (Zapata et al. 2002), (Fig. 5). If this interpretation is correct, part of the Upper Cretaceous Andean deformation had to affect the Agrio FTB. Field cross-cutting relationships on the Inner Sector of the Agrio FTB show that there are extrusive volcanic rocks of the Lower Eocene Cayanta Formation (44-49 Ma: Llambias & Rapela 1989) covering the previously deformed Cretaceous units of the Neuquen Basin (Fig. 3) (Repol et al. 2002). The Cayanta Formation is mostly preserved in a regional synclinorium (the Pichaihue syncline; Fig. 5) as lavas that overlie the Huitrin and Agrio formations. This fact, together with other field relationships, documents that during early Eocene times not only the Pichaihue syncline was already formed (it has clearly controlled the locus of the lava flows), but also there was more than 1000 m of uplift (the Rayoso and Neuquen groups were already eroded). Consequently, during the early Eocene, most of the major structures of the Agrio FTB were already formed. Finally, evidence of Late Cretaceous uplift is also found on the thrust-front structures of the Outer Sector of the Agrio FTB, where

the Puesto Burgos ignimbrites of early Eocene age (based on flora ages of the Nectandra patagonica sp.) unconformably overlie the lower units of the Neuquen Group (Zapata et al 2002). Synorogenic deposits of the Conglomerados de Tralahue (Ramos 1998) and He Rincon Bayo Formation (Zapata et al 2002) are associated with structures of both the Inner and Outer sectors of the Agrio FTB. These deposits are found on the limbs of the large anticlines showing growth strata relationships and, hence, demonstrate a Miocene episode of deformation (Ramos 1998; Repol et al 2002; Zapata et al 2002). Field relationships, together with seismic data, suggest that during this period some of the pre-existing structures were reactivated with no more than 500 m of uplift. The Agrio FTB structures were carried over the regional decollement (Fig. 5), as deformation was transferred toward the Chihuidos High and the Neuquen Basin platform structures (Figs 1 & 2). Unfortunately, the age of these units has been established using mammal fossil fauna recognized up to genus (not to species) such as Notoungulate (Repol et al 2002; Zapata et al 2002), and therefore the age range is somewhat broad. Regional correlations suggest that the upper limit for this deformation should not exceed 12 Ma (see discussion below). Main Andean cordillera Stratigraphy of the main corillera Mesozoic rocks. Jurassic rocks along the axial zone of the cordillera have formed part of an intra-arc basin, isolated in terms of superficial exposure from the western Neuquen Basin which constituted a retro-arc embayment at these latitudes. These rocks are upper Pliensbachian-middle Callovian and probably up to Kimmeridgian (Suarez & Emparan 1997), and comprise submarine clastic carbonate and andesitic-dacitic rocks. At the top of the sequence conditions became subaerial for the whole intra-arc axis and retro-arc area. Early Cenozoic volcanic rocks. The intrusive subvolcanic and volcanic facies associated with the Cura Mallin Formation of late OligoceneEarly Miocene age define an ancient volcanic belt in Chile (25-15 Ma) (Suarez & Emparan 1995, 1997; Burns & Jordan 1999; Jordan et al 2001; Radic et al. 2002). The corresponding arc front was located in the Coastal Cordillera (Munoz et al 2000; Stern et al 2000).

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Equivalent rocks are located in the Huantraico area, more than 300 km from the trench, where slab-related volcanics were dated as 20-18 Ma (Fig. IB, C) (Kay 2002). During Oligocene-Miocene times a regional episode of sedimentation caused by regional extensional deformation affected the Southern Central Andes (Fig. 6) (Cazau et al 1987; Herve et al 1995; Charrier et al 1996; Spalletti & Dalla Salda 1996; McDonough et al 1997; Godoy et al 1999; Munoz & Araneda 2000; Rivera & Cembrano 2000). Two units have been distinguished in the main cordillera: the Cura Mallin Formation (Niemeyer & Munoz 1983) that represents most of the intra-arc stratigraphic column; and the Mitrauquen Formation capping the sequence (Fig. 3) (Suarez & Emparan 1997). The Cura Mallin Formation is a volcanic-sedimentary unit, deposited in lacustrine and fluvial environments, dated between 20 and 11 Ma (Suarez & Emparan 1995; Jordan et al 2001; Radic et al 2002). The Mitrauquen Formation lies over the former and is composed of ignimbrites and upwards-coarsening fluvial conglomerates, with ages between 10 and 8 Ma (Suarez & Emparan 1997). The Cura Mallin Basin (Fig. 6) formed diachronously during the late Oligocene-Early Miocene north of 37°30/S (Jordan et al 2001), and during the Middle Miocene in the Lonquimay area south of 38°S (Fig. 4) (Suarez & Emparan 1995). The composition of the sediments accumulated in the basin is also variable between depocentres and is dependent on its regional position within the basin (Suarez & Emparan 1995; Jordan et al 2001; Radic et al 2002). The depocentre located north of 37°30/S has a predominance of sedimentary non-marine facies to the east, while to the west these are replaced by thick volcanic successions (Rovere et al 2000). However, in the southern depocentre, the sedimentary facies are located next to the western limit of the basin (Suarez & Emparan 1995). Based on this change in polarity a major transfer zone is located between these two areas, under the Quaternary Mandolegiie chain around 37°45/-38°S (Figs 6 & 7) (Melnick et al 2002; Radic et al 2002). The Cura Mallin Formation is either conformably overlain by the 10-8 Ma Mitrauquen Formation, or separated by an unconformity (Suarez & Emparan 1997). These deposits correspond with alluvial fans accumulated in intermontane basins, being proximal equivalents in the main cordillera of a foreland basin, equivalent to the Tralahue and Rincon Bayo Formations in the Inner Sector of the Agrio FTB.

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The Cura Mallin Basin is bounded by the Patagonian Batholith to the west and the western Mesozoic outcrops of the Neuquen Basin and the Cordillera del Viento (Inner Sector of the Agrio FTB) to the east (Figs IB, C&6). Its western basement is formed by Jurassic-Cretaceous granitoids corresponding to the Northern Patagonian Batholith. The eastern basement is Carboniferous-Permian volcanic and sedimentary rocks, Lower Jurassic sequences, and volcaniclastic and intrusives of 57-50 Ma (Fig. 3) (Ramos & Nullo 1993; Rovere 1993). Late Cenozoic volcanic rocks. Three different volcanic pulses with distinct arc fronts can be distinguished during the Late Cenozoic. (1) The oldest one corresponds to the Cola de Zorro Formation (Vergara & Munoz 1982), which is formed by volcanic breccias and lavas of basaltic-dacitic composition, minor sedimentary facies and related intrusive bodies (Fig. 3) (Vergara & Munoz 1982; Niemeyer & Munoz 1983). Its age is constrained between 5 and 3.5 Ma, and its maximum thickness is around 1900 m, although in most cases it varies between 100 and 1000 m. The regional distribution of this unit constitutes a band of northsouth development between 70°45/W and 71°15'W, and with a latitudinal range of 36°39°S (Fig. 6) (Niemeyer & Munoz 1983; Suarez & Emparan 1997; Linares et al 1999; Rovere et al 2000). (2) A relatively stable arc position corresponds to Upper Pliocene-Lower Pleistocene (2-0.5 Ma) volcanics that constitute a narrow belt with its front at 72°W (Lara et al 2001). According to these authors the Upper Pliocene volcanic width of the belt and its front are more or less coincident with the Lower Pliocene one. The width drastically diminished during the Late Pleistocene-Holocene, and attained a third stable arc position coincident with (3) the present volcanic front. The Cola de Zorro Formation covers and, consequently post-dates, the Late Miocene deformation (Fig. 8). This unconformity is of regional importance, existing at both sides of the Andes at these latitudes. Description of the structure Palaeogene structure - Cura Mallin Basin. The extensional nature of this basin has been interpreted from 2D seismic lines along the eastern edge of the basin (Burns & Jordan 1999; Jordan et al 2001). Two sectors (northern and southern) of the Palaeogene basin have been separated, based on sediment thickness, structure and age

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Fig. 6. Relationship between regional Mesozoic structure and the development of intra-arc and inner retro-arc extensional structures along the inner sectors of the Neuquen FTB. Modified from Manceda & Figueroa 1995; Vergara et al., \991a-c, Ramos 1998; Lavenu & Cembrano, 1999; Radic et al. 2002. Note the similarities between the Jurassic and Palaeogene rifting geometries.

(Radic et al. 2002): (1) a northern sector (24.622.8 Ma) is mainly located along the eastern slope of the Andean cordillera, where a thick sequence, around 2800 m, is exposed by a

W-dipping inverted extensional fault (Jordan et al 2001); and (2) a southern sector (17.5-13 Ma) 2600m thick (Suarez & Emparan 1995) (Fig. 7). Consequently, the Cura Mallm Basin

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Fig. 7. Bouger anomalies beneath the inner sectors of the Agrio FTB, and segmentation of the Palaeogene Cura Mallin Basin by the Mandolegiie transfer zone, similar to the geometries proposed for the Lower Jurassic rift by Manceda & Figueroa (1995) immediately to the north (see Fig. 6). Note the parallelism between the Lower Jurassic extensional structures corresponding to the Tres Chorros extensional system (Vergani et al. 1995; Ramos 1998) and the Neogene structure developed along the volcanic arc.

did not behave in a homogeneous way and, moreover, the corresponding polarity also alternated with time (Radic et al. 2002). The fold and thrust belt that closed the basin during the Late Miocene was controlled by polarity changes and transfer zones in the upper Oligocene rift architecture (Jordan et al. 2001; Melnick et al 2002; Radic et al 2002) (Figs 6-8). The Rio Picunleo constituted an east-west transfer zone during the formation of the fold and thrust belt at these latitudes (Fig. 8). To the north, an east-verging structure is exposed, meanwhile to the south between this river and the Las Damas valley the thrusts are dipping to the east (Fig. 9). Therefore,

based on the extensional structure that has controlled the basin, the Rio Picunleo would have been a transfer zone corresponding to extensional panels, subsequently inverted during the Late Miocene. Neogene—Quaternary structure. The Cola de Zorro Formation displays important changes in thickness, from more than 1000 m locally up to 1900 in the Laguna de la Laja area (Niemeyer & Mufioz 1983), to less than 200 m in the Sierra de Trocoman and other regions on the eastern side of the Andes at these latitudes.

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Fig. 8. (A) 3D shaded relief image corresponding to (B). (B) Block diagram showing alternating panels in the Upper Miocene compressive structure as a function of the Palaeogene rifting architecture. Note the superposition of the Lower Pliocene depocentres using transfer faults of Miocene age. (C) Internal unconformities in the Lower Pliocene sequences associated with synextensional accumulation.

The Sierra de Trocoman acted as a structural high during the Early Pliocene (Figs 8 & 9). The east-west Rio Picunleo and the east-west Las Damas valley flank this structural high and limit the two Lower Pliocene depocentres. Therefore, the Caldera del Agrio zone has been the locus of the Lower Pliocene accumulations south of the Las Damas valley, where the Cola

de Zorro unit dips to the north against a normal scarp along which the upper Palaeogene sequence is exposed. The unconformity between the Miocene and Lower Pliocene sequences is dismembered through the eastern slope of the Neuquen Andes, by the Neogene extensional structure. Particularly, the main N-S branch of the

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Fig. 9. (A) Shaded relief image, where the main structural features of the inner sectors eastern side of the main cordillera and Loncopue trough are represented. (B) 3D shaded relief image corresponding to (C). (C) Inner sectors of the Neuquen fold and thrust belt between 38° 15' and 37°30/S. Note that extensional structures are affecting the compressive structure at the northern part of the slide. These are interpreted as an Early Pliocene relaxation of the compressive wedge in response to subtle changes in subductive border conditions.

Rio Picunleo runs through an extensional fault that bounds the Picunleo -Renileuvu depocentre on its western side (Fig. 8). Therefore, two main sets of faults control the Lower Pliocene depocentres, one of N-S strike, which bounds the Upper Miocene fold and thrust belt on its eastern slope provoking the collapse of

the Palaeogene sequences, and the other of E-W strike defining a series of highs and lows aligned with the meridian (Figs 8 & 9). In addition, to the recognition of extensional faulting controlling the Lower Pliocene accumulations, evidence of synsedimentary deformation associated with the structures was also found in

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the area. Unconformities separate rotated packages of Lower Pliocene basal sequences from almost flat members corresponding to the same unit (Fig. 8). This situation is visible in the northern divide of the Picunleo valley, where a block has been rotated with N-S strike prior to the accumulation of the uppermost stage of the Cola de Zorro Formation (Fig. 8). To the south, along the northern side of the Caldera del Agrio, other synextensional faults have been identified in the Trolope valley (Fig. 9). Therefore, extensional activity is documented from the initial stages of the Early Pliocene, when the fold and thrust belt seems to have delayed its progression to the foreland and its inner sector collapsed forming a narrow intra-arc basin (Figs 6 & 8). A long intra-arc feature, the dextral strike-slip Liquine-Ofqui Fault Zone (LOFZ), active since at least the Late Cretaceous (Fig. 1) (Lavenu & Cembrano 1999; Cembrano et al 2000), ends in a more complex structural array at 38°S (Melnick et al 2003) (Figs 1 & 9). In this area, a series of NNW- to NW-striking faults, of which the BioBio left-lateral fault and the Chilian lineament are the main features (Fig. 2) (Ramos 1977; Dalla Salda & Franzese 1987; Mufioz & Stern 1988; Melnick et al 2002), are related to regional upper Palaeogene fractures, developed in the Southern Central Andes segment of the Andean chain between 36° and 40°S (Vergara et al \991a-c; Rivera & Cembrano 2000; Jordan et al 2001; Radic et al 2002). The area under extension coincides with the western Tres Chorros structural system inferred beneath the Cenozoic cover based on the main Mesozoic faults to the west (Fig. 7). Many Pliocene-Quaternary extensional faults recognized in the basement of the main polygenetic volcanoes, as the edges of the Caldera del Agrio (Fig. 9), coincide with the projections of these faults (Vergani et al 1995; Ramos 1998) to the arc platform (Fig. 7). The Liquine-Ofqui Fault Zone is a relatively continuous feature from 40° to 46°S, which runs through the Main Cordillera and the Quaternary volcanic arc (Lavenu & Cembrano 1999). The fault zone changes to a series of en echelon NNE to NE-trending faults between 38° and 39°S (Fig. 1). These faults have recently been recognized as normal faults related to the late Oligocene-Early Miocene extension, controlling the thickness of units of corresponding age (Radic et al 2002). Therefore, neotectonics associated with the volcanic front shows control by Palaeogene structures, as well as Mesozoic ones. The volcanic lineament Callaqui-CopahueCordillera de Mandolegiie (CCM) is the longest

volcanic fissure in the Cordillera Patagonica, reaching 80 km, and is transversally oriented to the volcanic-arc front (Fig. 7). This feature is formed by a series of coalescent polygenetic and monogenetic centres, which from west to east are: the Callaqui and Copahue volcanoes of 1-0 Ma; the Coladas de Fondo de Valle monogenetic field of 1.6-0.8 Ma; the Bayo and Trolope dome complexes of 0.6 Ma; and the Trolon eroded strato-volcano of 0.6 Ma (Pesce 1989; Linares et al 1999), together with an extensive monogenetic basaltic field that is part of the northern extent of the extensional Loncopue trough (Ramos & Folguera 19990, b). This belt lies in the middle of the rhomboedric Damas-Chaquilvm structure and crosses the Caldera del Agrio through its mid part (Fig. 9) (Melnick et al 2002). Finally, the Damas-Chaquilvm structure coincides with a major left-lateral transfer zone in the Cura Mallin Basin (Melnick et al 2002), forming part of the Quaternary volcanic basement, separating two diachronous sub-basins (Radic et al 2002). This observation supports the structural control exerted by the Palaeogene structure in the neotectonics and arc-front geometry (Fig. 7).

Uplift and temporal constraints Apatite fission-track ages from the main cordillera, immediately south of the studied sector (40°S), range between 4.3 and 1.3 Ma indicating a very young period of exhumation in the inner sectors of the fold and thrust belt (Fig. 4) (Grafe et al 2002). This uplift is out of sequence with respect to the rest of the CretaceousMiocene Agrio FTB. However, around 38°30/S, corresponding to the inner sector of the presently analysed section, the apatite fission-track ages fall within the Eocene, without any younger imprints (Fig. 4) (Grafe et al 2002).

Tectonic evolution of the Andean FTB Fission-track data together with Ar/Ar ages are conclusive in relation to the beginning of deformation of the Andes of the Neuquen Basin. Local angular unconformities, particularly at the northern and southern edges of the basin, reveal that retro-arc deformation was an active process during the Jurassic and Early Cretaceous (Vergani et al 1995). Such observations are consistent with the existence of a subduction zone and related magmatic arc (Kay et al 1989). However, the formation of the Andean FTB seems to have been delayed until the Late Cretaceous (Fig. 4) (Ramos 1977; Zapata et al 1999;

NEUQUEN FOLD AND THRUST BELT

Kay 2001). The Andean FTB was then characterized by the inversion of the pre-existing Jurassic extensional structures (Vergani et al 1995; Ramos 1998), now preserved in the Agrio FTB, associated with limited thin-skinned deformation focused at the basement block boundaries (Figs 5 & 9). The synorogenic-related sediments either are associated with the base of the Neuquen Group (Ramos 1977) or at least to the Cretaceous Malargiie Group deposits (Fig. 3) (Barrio 1990; Zapata et al. 2002). Deformation of the intra-arc region occurred at least until the Early Eocene, based on apatite fission-track data, indicating that the main Cordillera suffered an episode of uplift (Fig. 4) (Grafe et al 2002) related to the inversion of Lower Jurassic structures of a pre-existing intra-arc basin (De la Cruz & Suarez 1997). During this time, Lower Eocene intrusives were emplaced in the upper crust in previously deformed Upper Cretaceous structures. Early Oligocene-Early Miocene was a time of orogenic relaxation, which affected the core of the Andes with the opening of the Cura Mallin Basin. The extensional deformation was restricted to the west of the Cordillera del Viento (i.e. west of the Inner Sector of the Agrio FTB) (Figs 6 & 7) occupying part of the volcanic arc. The sedimentation in the intra-arc area was non-marine, although the Pacific Ocean occupied an important sector of the forearc area (McDonough et al 1997; Vergara et al I991a-c). The Late Miocene was, again, a period of contractional deformation. The Oligocene-Miocene intra-arc basin was closed due to minor reactivation of Mesozoic structures and localized inversion of intra-continental extensional faults of the Cura Mallin Basin (Fig. 8). The closure of the Cura Mallin Basin structures, as they were transported toward the foreland region, produced the reactivation of the Late Cretaceous Agrio FTB structures. The deformation front migrated eastwards toward the Chihuidos High and Platform areas of the Neuquen Basin, causing tectonic inversion of the pre-existing Jurassic half-grabens (Fig. 10). During the Early Pliocene, once again, tectonic relaxation affected a restricted latitudinal band between 36° and 39°S at the inner sectors of the Andean FTB, forming a narrow rift locally coincident with the previous Tertiary intra-arc basin (Figs 6 & 7). The corresponding normal faults displaced and buried the compressive structure formed during the Late Miocene, particularly at the eastern slope of the main Andes today (Fig. 10). Late Pliocene-early Quaternary times south of 37°30'S are marked by inner collapse of the fold and thrust structure

51

related to the northern prolongation of the dextral strike-slip Liquine-Ofqui Fault Zone (Ramos & Folguera 1999a, b; Melnick et al 2003). Fission track-data (Grafe et al 2002) reveal that, in contrast to the south of 39°S, at 38°30'S the main cordillera remained static during the Neogene, while to the south around 40°S out-of-sequence thrusting was related to a new period of fold and thrust development that is still active, as revealed by seismicity (Fig. 2) (Barrientos & Acevedo 1992; Bohm et al 2002).

Discussion The tectonic evolution of the Andes within the Neuquen Basin has experienced recurrent episodes of compression and extension that affected different components of the orogen. Each event has overprinted the previous one and, therefore, it is not likely that the available data allow a complete reconstruction of the history of the orogen. In addition, this study shows that all of the different structural and tectonic styles of deformation have been strongly controlled by the pre-existing extensional Jurassic features, mostly half-graben rift geometries driven by deep crustal discontinuities, associated with the opening of the Neuquen Basin. The first-order crustal features that define the deformation styles are the crustal discontinuities that may be a consequence of Late ProterozoicEarly Palaeozoic terrane accretion (Ramos 1989). Seismic, borehole, gravity and geochemistry data suggest that the depocentres of the Jurassic rift of the Neuquen Basin are concentrated on the upper plate of these crustal discontinuities (Zapata 1997). Following the simple-shear asymmetric extensional model (Wernicke & Burchfield 1982), the polarity of the crustal discontinuity should be directed towards the area of maximum subsidence, parallel to the basin half-graben polarity. Some of these first-order crustal discontinuities are located at the NE flank of the Neuquen Basin beneath the Auca Mahuida Volcano, with NE-directed polarity (Fig. 10) (Zapata 1997), and at the basin depocentre beneath the Chihuidos High (Ramas 2002; Zapata et al 2002), with E-directed polarity, etc. (Fig. 10). One of the most important crustal discontinuities is the one that lies immediately to the west of the Cordillera del Viento (Fig. 10) at the Inner Sector of the Agrio FTB. Borehole and seismic data show that the Jurassic synrift depocentres become progressively deeper toward the west as they lie closer to the crustal discontinuity,

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Fig. 10. (A) Shaded relief image where, cross-section position corresponding to (C) is displayed. (B) Exaggerated topography along (C), to visualize Loncopue-active extensional retro-arc zone at the transition between main cordillera and Agrio FTB. (C) Cross-section near 38°S from the Quaternary volcanic axis to the retro-arc area.

showing that the amount of subsidence was higher in that direction. Consequently, the crustal discontinuity associated with the asymmetric Jurassic rift should be westward-directed (Zapata et al 1999) (Fig. 10). According to this model, the asymmetry of the Cura Mallin Oligocene basin depocentres (Jordan et al. 2001) may be related to such crustal discontinuity (Fig. 10). Gravity data show that there is another major westward-directed crustal discontinuity beneath the actual volcanic arc region between 31°30' and 38°S (Fig. 7). This discontinuity may have controlled the change in polarity of the Cura Mallin and younger extensional basins recognized by previous workers in the area (Melnick et al. 2002; Radic et al. 2002). Finally, the second-order crustal discontinuities are those related to the geometry of the Jurassic half-graben. These pre-existing

structures have controlled the vergence and geometry of the Andean structures of the Agrio FTB, as well as the local depocentres and transfer zones of the Cura Mallin and younger basins, as it was described before (Figs 6 & 7).

Conclusions The Andean building processes determined from the presented data contain the following elements. •

First-order crustal discontinuities have controlled the deformational styles. A Late Cretaceous compressional event has been preserved at the present volcanic arc and Agrio FTB areas. This event produced a significant amount of uplift, particularly concentrated on the Cordillera del Viento (Inner Sector of the Agrio FTB), as a reactivation of the

NEUQUEN FOLD AND THRUST BELT

•

•

•

pre-existing rift geometry controlled by a westward-directed crustal discontinuity (Fig. 10). This event was also extended towards the foreland area on the Outer Sector of the Agrio. Unfortunately, there are not enough data preserved to reconstruct the uplift history for the present volcanic arc. These compressional events may have lasted until the earliest Eocene. During the latest Oligocene - earliest Miocene, as a consequence of a change in the geodynamic conditions (Jordan et al. 2001), a regional extensional event formed the Cura Mallin intra-arc basin. This extensional event was confined to a region to the west of the Cordillera del Viento where the pre-existing westward-directed crustal discontinuity was reactivated. No extensional movements have been recorded further to the east, toward the foreland basin (Fig. 4). For the Late Miocene, compressional movements closed the Cura Mallin Basin and reactivated the Agrio FTB structures. The shortening was transferred toward the foreland using the inherited Jurassic detachment (reactivated as a decollement) (Fig. 10). Early Pliocene - Quaternary has been a period of relaxation of the inner sectors of the Agrio Fold and Thrust Belt, probably reactivating a Jurassic deep detachment, as suggested at a local scale by (a) parallelism between Lower Jurassic structures of the Tres Chorros extensional system and Neogene structures; and at a broad scale by (b) similarities between Tertiary rift geometry and Jurassic rift geometry exposed north of 36°S (Fig. 6). In other cases, Pliocene-Quaternary extensional faults have cut the Upper Miocene compressional structure without tracing any particular pattern (Fig. 8).

The structural evolution of the Neuquen Basin Andean orogeny is characterized by compressional events affecting the region from the present volcanic arc through to the foreland. However, the extensional events have been restricted to the major crustal boundaries of the present volcanic arc and high cordillera (plateau?) regions, with minor propagation toward the foreland area. The authors acknowledge Repsol-YPF, LAM Direction for authorization to publish borehole, seismic and fission-track data. The authors specially acknowledge K.C. Morley and H. Welsink for their thoughtful review and constructive comments that improved the quality of this work. This study has been partially made with funding from PIP 4162 (CONICET) and PICT 059 for

53

the 'Study of the segment of normal subduction —33° to 38°S'.

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Andes (37° to 37°15'S-71°W). Abstracts of the Japan Volcanological Society, 2, 107. ROVERE, E. 1998. Volcanisme Jurâsico, Paleogeno y Neogeno en el Noroeste del Neuquén, Argentina. X° Congreso Latinoamericano de Geologia y VF Congreso Nacional de Geologia Economica, Actas, 1, 144-149. ROVERE, E., LEANZA, H., HUGO, C., CASSELLI, A., TOURN, S. & FOLGUERA, A. 2000. Hoja geologica Andacollo, Provincia de Neuquen, 1:250 000. Servicio Nacional de Geologia y Mineria Argentino, Buenos Aires. SPALLETTI, L.A. & DALLA SALDA, L. 1996. A pull apart volcanic related Tertiary basin, an example from the Patagonian Andes. Journal of South American Earth Sciences, 9, 197—206. STERN, C., Munoz, J., TRONCOSO, R., DUHART, P., CRIGNOLA, P. & FARMER, G. 2000. Tectonic setting of the Mid-Tertiary Coastal Magmatic belt in South Central Chile: an extensional event related to Late Oligocène changes in plate convergence rate and subduction geometry. IX° Congreso Geologico Chileno, (Puerto Varas), Actas, 2, International Symposium 3, 693—696. SUÂREZ, M. & EMPARÂN, C. 1995. The stratigraphy, geochronology and paleophysiography of a Miocene fresh-water interarc basin, southern Chile. Journal of South American Earth Sciences, 8, 17-31. SUÂREZ, M. & EMPARÂN, C. 1997. Hoja Curacautm. Regiones de la Araucania y del Bio Bio. Carta Geolôgica de Chile, 1:250 000. Servicio Nacional de Geologia y Mineria de Chile, Santiago, 71, 105. THOMSON, S. 2002. Late Cenozoic geomorphic and tectonic evolution of the Patagonian Andes between latitudes 42°S and 46°S: An appraisal based on fission-track results from the transpressional intra-arc Liquine-Ofqui fault zone. Geological Society of America, Bulletin, 114, 1159-1173. VERGANI, G., TANKARD, A., BELOTTI, H. & WELSINK, H. 1995. Tectonic evolution and paleogeography of the Neuquén Basin Argentina. In: TANKARD, A.J., SUÂREZ, R. & WELSINK, H.J. (eds), Petroleum

Basins of South America. AAPG Memoirs, 62, 383-402. VERGARA, M. & MUNOZ, J. 1982. La Formacion Cola de Zorro en la Alta cordillera Andina Chilena (36°-39°S), sus caracteristicas petrogrâficas y petrologicas: Una revision. Revista geologica de Chile, 17, 31 -46. VERGARA, M., LOPEZ ESCOBAR, L. & HICKEYVARGAS, R. 1997^. Geoquimica de las rocas volcânicas miocenas de la cuenca intermontana de Parral-Nuble. VIII0 Congreso Geologico Chileno (Antofagasta), Actas, 2, 1570-1573, VERGARA, M., MORAGA, J. & ZENTILLI, M. \991b. Evolucion termotectonica de la cuenca terciaria entre Parral y Chilian: Anâlisis por trazas de fision en apatitas. VIII0 Congreso Geologico Chileno (Antofagasta), Actas, 2, 1574-1578. VERGARA, M., PUGA, E., MORATA, D., BECCAR, L, DIAZ DE FEDERICO, A. & FONSECA, E. 1997c. Mineral chemistry of the Oligocène-Miocene volcanism from Linares to Parral, Andean Precordillera. VIII0 Congreso Geologico Chileno (Antofagasta), Actas, 2, 1579-1583. WERNICKE, B. & BURCHFIELD, B. 1982. Models of extensional tectonics. Journal of Structural Geology, 4, 105-115. ZAPATA, T. 1997. La estructura cortical del volcan Auca Mahuida. YPFSA internal Report, Bueros Aires. ZAPATA, T., BRISSÔN, I. & DZELALIJA, F. 1999. The role of basement in the Andean fold and thrust belt of the Neuquén Basin. In: Thrust Tectonics. University of London, London, Abstracts, 122124. ZAPATA, T., CÔRSICO, S., DZELAJICA, F. & ZAMORA, G. 2002. La faja plegada y corrida del Agrio: Anâlisis estructural y su relacion con los estratos terciarios de la Cuenca Neuquina Argentina. V° Congreso de Exploraciôn y Desarrollo de Hidrocarburos (Mar del Plata), Actas. IAPG, Buenos Aires, electronic files. ZOLLNER, W. & AMOS, A. 1973. Descripcion Geologica de la Hoja 32b, Chos Malal, 1:200000. Carta Geologico Economica de la Repubica Argentina, Buenos Aires, 143, 91.

Lower Cretaceous (Berriasian-Aptian) biostratigraphy of the Neuquén Basin M. BEATRIZ AGUIRRE-URRETA1, PETER F. RAWSON2, G. ANDREA CONCHEYRO1, PAUL R. BOWN2 & EDUARDO G. OTTONE1 l Laboratorio de Bioestratigrafia de Aha Resolution, Departamento de Ciencias Geologicas, Universidad de Buenos Aires, Pabellôn H, Ciudad Universitaria, CI428EHA, Buenos Aires, Argentina (e-mail: [email protected]) ^Department of Earth Sciences, University College London, Gower Street, London WC1E 6BT,UK Abstract: The Berriasian-Aptian succession in the Neuquén Basin is mainly marine in the lower part and non-marine in the upper portion. A detailed ammonite zonation is presented for the Berriasian-?Early Barremian interval. While some ammonite taxa are endemic, others are widely distributed and there are several levels where correlation can be suggested with the 'standard' stages and zones of the Tethyan Mediterranean area. Several nannofossil bioevents are recognized, and these provide evidence for correlation with Tethyan areas. Correlations suggested by both groups are reasonably consistent. Berriasian-Aptian palynomorphs include both terrestrial and marine forms. Several terrestrial assemblages can be recognized, but the marine forms are mainly long-ranging taxa, especially in the Agrio Formation.

The infill of the Neuquén Basin comprises more than 6000 m of marine and continental sedimentary rocks ranging in age from Late Triassic to Palaeogene. The Lower Cretaceous succession is represented by the Mendoza (part) and Rayoso groups. Sediments of Berriasian-?Early Barremian age are mostly marine and crop out extensively in the Neuquén and Mendoza provinces yielding rich marine faunas and floras. In Neuquén they form the upper part of the Vaca Muerta Formation, the Mulichinco Formation and the Agrio Formation of the Mendoza Group. Here ammonites are often abundant and well preserved. Further north, in southern and central Mendoza, the faciès are generally calcareous (Chachao and Agrio formations) and ammonites are less common and often flattened. Conversely, nannofossils are more abundant in the more calcareous faciès. Palynomorphs have been studied mainly in the Neuquén sections. During most of the Barremian up to the Aptian, sedimentation in the basin was characterized by evaporites and red continental deposits of the Huitrin and Rayoso-Ranquiles formations (Rayoso Group) (Fig. 1). In these rocks, only palynomorphs have been recovered. The research focused initially on establishing a detailed ammonite succession (see below), to

improve correlation of the marine beds both within the basin and with the succession in the Mediterranean area, where the standard stages were defined. Then the work expanded to embrace other biostratigraphically important and geographically widespread fossil groups in marine succession, of which nannofossils and palynomorphs (especially dinoflagellates) are potentially the most useful over the interval of time represented, when planktonic foraminifera were still rare. Many of the palynomorphs proved to be long-ranging forms of little value in long-distance correlation, but the nannofossils provided a valuable cross-check on the correlation indicated by the ammonites and thus on the application of the 'standard' stage names to the Neuquén Basin. This paper reviews and synthesizes the work completed to date, and also summarizes the palynology of the overlying non-marine beds of the Rayoso Group. AguirreUrreta and Rawson are responsible for the ammonite studies, Bown and Concheyro for the nannofossils, and Ottone for the palynomorphs.

Ammonite biostratigraphy The main descriptions of the ammonites of the Vaca Muerta Formation (Tithonian-Lower

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquén Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 57-81. 0305-8719/05/$15.00 © The Geological Society of London 2005.

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Fig. 1. Generalized stratigraphie columns of the Lower Cretaceous Mendoza and Rayoso groups in Mendoza and Central Neuquén.

Valanginian) are by Burckhardt (1903), Gerth (1925), Weaver (1931), Windhausen (1931), Leanza (1945) and Aguirre-Urreta & Alvarez (1999). Leanza (1945) concentrated on faunas from Mendoza and for the Berriasian-earliest Valanginian interval, and proposed three biozones, Argentinicems noduliferum, Spiticems damesi and lNeocomites' wichmanni, all of which are still used. Leanza (198la, b) and Riccardi (1984) listed the faunas of each biozone. With the exception of a study on Groebericeras (Aguirre-Urreta & Alvarez 1999) there has been little recent systematic work on the faunas of this age, but Riccardi (1988) and AguirreUrreta (1993) have both provided a general biostratigraphy with illustrations of the most characteristic ammonites. Conversely, over the last 12 years, the late Early Valanginian-latest

Hauterivian/Early Barremian faunas have been studied in detail, the authors spending several field seasons visiting numerous localities (Fig. 2) and collecting the rich ammonite faunas that occur in the top Vaca Muerta, Mulichinco and Agrio formations. This led to the publication of a much more detailed and accurate biostratigraphy than any previous one (Aguirre-Urreta & Rawson 1997). In that paper we noted that many of the cited taxa were in need of systematic revision. Since then we have collected much additional material and have revised the systematics of some of the key genera (Aguirre-Urreta 1998; AguirreUrreta & Rawson 1999a-c, 2001, 2002, 2003). Others still await revision. Hence, although we still recognize the same general succession of faunas, the names of some of the index taxa

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Fig. 2. Principal localities of the Neuquén Basin where recent studies on ammonites, palynomorphs and nannoplankton have been made.

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have had to be changed, resulting in changes to some subzonal names (Fig. 3). We have also been able to define some of the zonal/subzonal boundaries with greater precision, and to establish the precise levels at which some of the less well-known forms occur (e.g. Lissonia, Valanginites, Hoplitocrioceras). The zonal sequence that we now recognize is comparable in detail to zonations for both the 'standard' Mediterranean sequences and the Boreal ones of NW Europe. The Neuquén faunal succession is now known in more detail than that in any other South American or Gondwanan sequence of similar age, and thus serves as a model for comparison across the region. There are some ammonites in common with the Austral Basin of southern Patagonia (Olcostephanus, Chacantuceras, Crioceratites), while others (e.g. Lissonia) extend northwards from Neuquén to northern Chile, Peru and Colombia.

The zonation is shown in Figures 3 (partially), 8 and 9, and individual zones are discussed briefly below. Examples of each zonal/subzonal index fossil are illustrated in Figs 4-7. In general, the faunas in any one subzone are of low generic diversity, and diversity generally decreases upwards, so that most of the Hauterivian and ?Early Barremian assemblages are monogeneric. Zonal boundaries mark major faunal turnovers; some of the subzonal boundaries also do, while others reflect an evolutionary sequence within a particular taxonomic group. The Juras sic-Cretaceous boundary in the Neuquén Basin has traditionally been marked on biostratigraphic grounds, as there is a complete lithological continuity within the black shales of the Vaca Muerta Formation across the passage from one system to another. The boundary has habitually been placed between the Substeueroceras koeneni Zone (Upper Tithonian)

Aguirre-Urreta & Rawson 1997

This paper

BIOZONE Weavericeras vacaensis

SUB-BIOZONE

BIOZONE Weavericeras vacaensis

SUB-BIOZONE

Hoplitocrioceras gentilii

Hop. gentilii

Hoplitocrioceras gentilii

Hop. gentilii

Holcoptychites neuquensis

Hop. sp. nov. Olcostephanus (O.) leanzai Hoi. compressum

Holcoptychites neuquensis

Hoi. neuquensis

Neocomites wichmanni

Olcostephanus (O.) laticosta Hoi. agrioensis Hoi. neuquensis

Neocomites sp. Pseudofavrella 'Acanthodiscus' sp. angulatiformis Pseudofavrella angulatiformis O. (Lemurostephanus) sp. Olcostephanus Karakaschiceras (Olcostephanus) attenuatus atherstoni Olcostephanus (O.) atherstoni

Hop. giovinei

Pseudofavrella angulatiformis

Olcostephanus (Olcostephanus) atherstoni

Neocomites sp. Chacantuceras ornatum Pseudofavrella angulatiformis O. (Viluceras) permolestus Karakaschiceras attenuatus Olcostephanus (O.) atherstoni

Lissonia riveroi Neocomites wichmanni

Fig. 3. Early Valanginian-Early Hauterivian ammonite zones and subzones of the Neuquén Basin: comparison between the proposal of Aguirre-Urreta & Rawson (1997) and this paper.

LOWER CRETACEOUS BIOSTRATIGRAPHY and the Argentiniceras noduliferum Zone (Lower Berriasian). However, Riccardi et al (2000) recently placed most of the Substeueroceras koeneni Zone in the Berriasian without explanation. Howarth (1992) in his study on the Tithonian and Berriasian ammonites from Iraq produced a correlation chart, including two different zonal schemes for South America. The traditional scheme was after Jeletzky (1984) while the other one was from Zeiss (1986), and in the latter the Substeueroceras koeneni Zone represents a large portion of the Berriasian. As further work is necessary to solve these problems, which are out of the scope of this paper, the traditional zonation is provisionally followed here.

The Argentiniceras noduliferum Assemblage Biozone (Leanza 1945) The zone was proposed by Leanza (1945) for southern Mendoza and placed in the Lower Berriasian. This author studied the ammonites that were collected by Groeber in southern Mendoza, and there has been no further research on this fauna, except the study on Groebericeras by Aguirre-Urreta & Alvarez (1999). Species of Argentiniceras (Fig. 4c), Berriasella, 'Thurmannic eras', Frenguelliceras and Substeueroceras are other components of this zone.

The Spiticeras damesi Assemblage Biozone (Gerth 1921) This zone was proposed by Gerth (1921) who assigned it to the Valanginian. Later, Burckhardt (1930) transferred it to the Berriasian. Leanza (1945) studied the ammonites and placed the index species S. damesi (Fig. 4d) in the Upper Berriasian together with Cuyaniceras transgrediens. The diverse components of the fauna, including species of Neocosmoceras, Neocomites and 'Thurmanniceras', await detailed sampling and a modern systematic revision.

The Neocomites wichmanni Assemblage Biozone (Leanza 1945) Proposed for beds containing the index species (Fig. 4e, f) and 'Thurmannites pertransiens Sayn'. Aguirre-Urreta & Rawson (I999a) figured some forms from this zone as 'Thurmanniceras' but noted that the faunas needed thorough revision. At about the boundary between this and the overlying Lissonia riveroi Zone is a thin horizon with Valanginites

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argentinicus Leanza & Wiedmann; a single Olcostephanus sp. is probably from the Valanginites bed and represents the earliest Argentine record of this genus (Aguirre-Urreta & Rawson 1999a).

The Lissonia riveroi Local Range Biozone (Aguirre-Urreta & Rawson 1999a) The base of the zone is defined by the first appearance of the index species, which apparently evolved from late 'Thurmanniceras' and co-occurs with Acantholissonia gerthi (Weaver) in the highest part of the Vaca Muerta Formation in Neuquén (Aguirre-Urreta & Rawson I999a). Lissonia riveroi (Lisson) (Fig. 4a, b) is a widely distributed South American form, known from Chile, Peru and Colombia.

The Olcostephanus atherstoni Assemblage Biozone (Aguirre-Urreta & Rawson 1997) Proposed as a new name for the former zone of Olcostephanus curacoensis (named by Leanza 1945) because the original index species is regarded as a junior subjective synonym of O. atherstoni (Sharpe). The base of the zone is defined by the first appearance of Olcostephanus atherstoni (Fig. 5a), which is often within the upper part of the Mulichinco Formation. This marks a major faunal turnover, from neocomitid to olcostephanid ammonites. The zone is developed mainly in the lower part of the Pilmatué Member of the Agrio Formation. Three subzones are distinguished. At the base is the O. (O.) atherstoni Subzone (AguirreUrretta & Rawson 1997), which contains abundant examples of the index species, are both microconchs and large macroconchs up to 150 mm in diameter. The base of the overlying Karakaschiceras attenuatus Subzone (AguirreUrreta & Rawson 1997) is defined by the sudden appearance of abundant neocomitids (Karakaschiceras and Neohoploceras: AguirreUrreta 1998) (Fig. 5e). Olcostephanus atherstoni co-occurs in the lowest beds, then disappears. A group of unusually evolute olcostephanids eventually replaced the neocomitids to characterize the O. (Viluceras) permolestus Subzone. First proposed as the O. (Lemurostephanus) sp. Zone by Aguirre-Urreta & Rawson (1995) to accommodate Leanza's (1958) 'Simbirskites' fauna, it was renamed by Aguirre-Urreta & Rawson (1999c) following taxonomic revision of the characteristic fauna. The base of the subzone is defined by the first appearance of Viluceras (Fig. 5b). Evolute forms of Olcostephanus

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Fig. 4. Berriasian-Early Valanginian index ammonites from the Vaca Muerta Formation, (a) & (b) Lissonia riveroi (Lisson) (CPBA 17307). (c) Argentiniceras noduliferum (Steuer) (CPBA 17306). (d) Spiticeras damesi (Steuer) (CPBA 7606). (e) & (f) Neocomites wichmanni (DNGM 7260). CPBA, Repository of the University of Buenos Aires, Argentina. DNGM, Repository of the Geological Survey of Argentina.

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Fig. 5. Valanginian index ammonites from the Agrio Formation, (a) Olcostephanus (Olcostephanus) atherstoni (Sharpe) (CPBA 11481). (b) Olcostephanus (Viluceras) permolestus (Leanza) (CPBA 19149). (c) & (d) Pseudofavrella angulatiformis (Behrendsen) (CPBA 17308). (e) Karakaschiceras attenuates (Behrendsen) (CPBA 17309). (f) Chacantuceras ornatum Aguirre-Urreta & Rawson (CPBA 18380). (g) Neocomites sp. (CPBA 17310).

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sensu stricto form a minority element of the fauna (Aguirre-Urreta & Rawson 1999c, fig. 4a-d, g, h, k).

The Pseudofavrella angulatiformis Assemblage Biozone (Aguirre-Urreta & Rawson 1995) The name was proposed as a replacement for the 'Lyticoceras' pseudoregale Zone of Gerth (1925). 'Lyticoceras' pseudoregale is poorly known and often misinterpreted, while Pseudofavrella angulatiformis (Fig. 5c, d) is a characteristic and distinctive component of the lower beds. The base of the zone marks another major faunal turnover, from olcostephanid to neocomitid ammonites, with slight overlap in the lowest bed at some localities. Neocomitids characterize the whole zone. Three subzones were proposed by AguirreUrreta & Rawson (1997), those of P. angulatiformis, 'Acanthodiscus' sp. and 'Neocomites' sp. At the base, a Pseudofavrella- 'Besairieceras' fauna characterizes the P. angulatiformis Subzone. The fauna is widely distributed, although often flattened in dark shales. But at Pichaihue (Fig. 2) it is beautifully preserved and bed-by-bed collecting has shown that P. garatei Leanza & Wiedmann is the first species to appear. A few metres higher it is largely replaced by Pseudofavrella angulatiformis (Behrendsen) and then 'Besaireiceras' australe (Leanza & Wiedmann) appears a little higher. The fauna still awaits revision, but examples of all three species were described and figured by Leanza & Wiedmann (1980); their ' Acanthodiscus vacecki (Neumayr & Uhlig)' is an advanced growth stage of 'B. ' australe. The 'Acanthodiscus ' sp. Subzone is characterized by a very distinctive, strongly tuberculate neocomitid that we have since described as a new genus, Chacantuceras: as a result, the name of the subzone was changed to that of Chacantuceras ornatum (Fig. 5f) (AguirreUrreta & Rawson 19996). The fauna of the Neocomites sp. Subzone consists of involute, compressed neocomitids (Fig. 5g), some closely similar in lateral view to European forms of the Neocomites (Teschenites) pachydicranus group. The fauna awaits description: two specimens were figured by Aguirre-Urreta & Rawson (1997, fig. 7d-f) as Neocomites sp. nov. In the highest part of the subzone and in the lowest part of the overlying H. neuquensis Subzone rare Oosterella occur (Aguirre-Urreta & Rawson 1996, 2003).

The Holcoptychites neuquensis Assemblage Biozone (Gerth 1921, modified) As originally defined (Gerth 1921, p. 143), the zone extended from immediately above the beds with Neocomites (the top of Gerth's 'Acanthodiscus radiatus Zone') to immediately beneath the first appearance of Crioceratites (his 'Crioceras andinum Zone'). It thus embraces several distinct faunal horizons through a considerable thickness of sediment. Aguirre-Urreta & Rawson (1997) restricted it to include only those beds characterized by Holcoptychites, plus a thin horizon immediately above that contains Olcostephanus. They recognized three subzones, of H. neuquensis, H. compressum and Olcostephanus (O.) leanzai. Since then all the constituent faunas have been monographed (Aguirre-Urreta & Rawson 2002, 2003); as a result, two of the subzonal names have changed (Fig. 3). The base of the zone, and of the H. neuquensis Subzone, is marked by the first appearance of the genus Holcoptychites. The lowest forms (H. cf. recopei (Douvillé) and H. sp. nov.) are poorly preserved; higher in the subzone H. magdalenae (Douvillé) appears, then H. neuquensis (Douvillé) (Fig. 6b, c). The middle subzone is characterized by more compressed Holcoptychites originally assigned to H. compressum Leanza & Wiedmann. That species is now regarded as a junior subjective synonym of H. agrioensis (Weaver) (Fig. 6d) so the name of the subzone has had to be changed (Aguirre-Urreta & Rawson 2003). The base of the highest subzone is marked by the abrupt replacement of Holcoptychites by Olcostephanus - the fourth, and final, invasion of the basin by that genus. The fauna is one of the most widespread ones in the basin. The index species was initially identified as Olcostephanus leanzai (Giovine), but Aguirre-Urreta & Rawson (2002) showed that O. (O.) leanzai is a junior subjective synonym of O. (O.) laticosta (Gerth) (Fig. 6a) and thus changed the subzonal name to the laticosta Subzone. Apart from the index species, a single O. (O.) boesei (Riedel) is recorded, while in the upper part of the subzone O. (Jeannoticeras) agrioensis Aguirrre-Urreta & Rawson also occurs. Seven fragments of an indeterminate neocomitid are known from the highest 2 m of the zone (Aguirre-Urreta & Rawson 2002) and may be the predecessor of the Hoplitocrioceras of the zone above. Aguirre-Urreta & Rawson (2003) have proposed Agua de La Mula as the standard reference section for the neuquensis Zone and its three subzones.

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Fig. 6. Hauterivian index ammonites from the Agrio Formation, (a) Olcostephanus (Olcostephanus) laticosta (Gerth) (CPBA 17311). (b) & (c) Holcoptychites neuquensis (Douville) (holotype EM 2001). (d) Holcoptychites agrioensis (Weaver) (CPBA 20011.2). EM, Repository of the University Claude Bernard, Lyon, France.

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The Hoplitocrioceras gentilii Assemblage Biozone (Aguirre-Urreta & Rawson 1997) Proposed for the middle part of the Holcoptychites neuquensis Zone of previous authors, the base of the zone is marked by another turnover at family level, the neocomitid genus Hoplitocrioceras replacing Olcostephanus. Hoplitocrioceras may be endemic to the basin. The gentilii Zone was divided into the Hoplitocrioceras sp. nov. and H. gentilii (Fig. 7a) subzones by Aguirre-Urreta & Rawson (1997), reflecting the evolution from more inflated, strongly tuberculate Hoplitocrioceras to more compressed forms. The fauna has since been monographed and the former 'H. sp. nov.' named as a new species, H. giovinei (Fig. 7b, c), by Aguirre-Urreta & Rawson (2001). As a result, the name of the subzone was also revised (see Fig. 3). Agua de La Mula was proposed as the standard reference section for the gentilii Zone and its two subzones. The only other ammonites recorded from the Hoplitocrioceras beds are two specimens of Olcostephanus (Olcostephanus) variegatus (Paquier) from the upper part of the giovinei Subzone (Aguirre-Urreta & Rawson 2002, fig. 5a-c). The Weavericeras vacaensis Assemblage Biozone (Aguirre-Urreta & Rawson 1997) Hoplitocrioceras is abruptly replaced by the desmoceratid(?) genus Weavericeras, which extends through the highest part of the Pilmatué Member of the Agrio Formation, up to the base of the Avilé Member. The two genera overlap in the lowest bed in the vacaensis Zone. Weavericeras (Fig. 7g) shows some variation in degree of inflation, but we have yet to investigate whether this has any stratigraphie significance. The genus was proposed by Leanza & Wiedmann (1980) who figured a single specimen and listed previous illustrations by Weaver (1931) and Giovine (1950). The Spitidiscus riccardii Assemblage Biozone (Aguirre-Urreta, Gutierrez Pleimling & Leanza 1993) The zone was proposed for blue-black shales with Spitidiscus at the base of the Agua de la Mula Member, immediately overlying the Avilé Member. There appear to be at least two species, S. riccardii (Leanza & Wiedmann 1992; Aguirre-Urreta 1995) (Fig. 7e, f) and an undescribed form (Aguirre-Urreta & Rawson

1997, p. 456, fig. 6e-g), but despite extensive searches we have yet to find both in the same section and therefore remain uncertain of the order of occurrence. The shales are widespread, but at many localities Spitidiscus is either absent or flattened in the shale and easily missed. Elsewhere solid body chambers are common, sometimes with some phragmacone. The Crioceratites schlagintweiti Consecutive-Range Biozone (Aguirre-Urreta & Rawson 1993) Proposed for the lower part of the former Crioceratites andinum Zone to include those beds immediately above the riccardii Zone in which the earliest Crioceratitidae appear, Crioceratites schlagintweiti (Giovine) (holotype refigured by Riccardi 1988, plate 8, figs 1 & 2; Aguirre-Urreta 1993, plate 4, fig. 3) (Fig. 7i) and C. apricus (Giovine) (holotype refigured by Riccardi 1988, plate 7, figs 3-5; AguirreUrreta 1993, plate 3, fig. 7). The fauna awaits revision. The Crioceratites diamantensis Consecutive-Range Biozone (Aguirre-Urreta & Rawson 1993) Proposed for the upper part of the former Crioceratites andinus Zone. The first appearance of the index species C. diamantensis (Fig. 7h) defines the base of the zone. Crioceratites andinus (Gerth) also occurs: the holotypes of both species were refigured by Aguirre Urreta (1993, plate 3, fig. 8, plate 4, fig. 1). Other specific names are also available, and the whole fauna awaits revision. But in comparison with the underlying forms, these later Crioceratites are more compressed and involute, showing a tendency to recoil. At Pichaihue some Crioceratites occur in the highest beds of this zone, but are different from the diamantensis—andinus group and look more like European Barremian forms. The Paraspiticeras groeberi Local Range Biozone (Aguirre-Urreta & Rawson 1993) The base of the zone is marked by the first appearance of Paraspiticeras (Fig. 7d), a few metres above the Crioceratites-bQaring beds. Only the index species is known. At Mina San Eduardo, Pichaihue, Puesto Ponce and Loma Tilhué, only a few fragments

LOWER CRETACEOUS BIOSTRATIGRAPHY

Fig. 7. Hauterivian and Early Barremian (?) index ammonites from the Agrio Formation, (a) Hoplitocrioceras gentilii Giovine (CPBA 19187). (b) & (c) Hoplitocrioceras giovinei Aguirre-Urreta & Rawson (CPBA 19234). (d) Paraspiticeras groeberi Aguirre-Urreta & Rawson (CPBA 17312). (e) & (f) Spitidiscus riccardii Leanza & Wiedmann (CPBA 17026). (g) Weavericeras vacaensis (Weaver) (CPBA 17313). (h) Crioceratites diamantensis (Germ) (MLP 21386). (i) Crioceratites schlagintweiti (Giovine) (holotype CPBA 5147). MLP, Repository of the La Plata Museum of Natural Sciences, La Plata, Argentina.

67

68

M. B. AGUIRRE-URRETA ETAL

of an open coiled heteromorph ammonoid have been recovered from the highest beds of the Agua de La Mula Member, above Paraspiticeras. The scarcity of the fossil material and the uncertainty of its systematic assignment prevent the recognition of a new zone. Correlation with the 'standard' succession of the Mediterranean region The Neuquen ammonite fauna consists of an alternation of endemic and more widely distributed genera. The latter include Spiticeras, Neocosmoceras, Neocomites, Karakaschiceras, Neohoploceras, Valanginites, Olcostephanus, Spitidiscus, Crioceratites and Paraspiticeras, all of which occur also in the Mediterranean area, where the 'standard' Lower Cretaceous stages and ammonite zones were defined (see Rawson et al 1996; Hoedemaeker et al 2003). But most of these genera were quite long ranging and embraced numerous species, very few of which are common to both areas. Hence, only a limited number of levels can be correlated with confidence. The preliminary correlations of Aguirre-Urreta & Rawson (1997) and Rawson (1999) are updated here in the light of much new stratigraphic and taxonomic work, and are summarized in Figures 8 and 9. In the Berriasian there are some genera in common between the Neuquen Basin and the Mediterranean region, as Spiticeras and Neocosmoceras are both well known from the latter area (Simonescu 1899; Djanelidze 1922; Mazenot 1939). However, it is not possible at the moment to compare them at a specific level due to the lack of modern studies of the Argentine faunas. Argentiniceras and Groebericeras are also known from the Mediterranean region but with more restricted occurrences (Pomel 1889; Mazenot 1939; Hoedemaker 1982). There is a good correlation in the middle part of the Valanginian, at two successive levels. The atherstoni Subzone of the Argentine succession can be correlated with the middle part of the campylotoxus Zone where O. guebhardi, a possible synonym of O. atherstoni, has its acme. Just above, the Argentine Karakaschiceras/Neohoploceras faunas of the attenuatus Subzone (Aguirre-Urreta 1998) are so close to those of the Mediterranean area that they can be correlated confidently with the biassalensis-peregrinus subzones of the French successions. There is then an extensive succession of faunas from the permolestus-compressum subzones where correlation is very tentative. In the

permolestus Subzone Viluceras, an endemic subgenus of Olcostephanus, is accompanied by occasional evolute Olcostephanus s.s. that provide a tentative link with the nicklesi Subzone of the Mediterranean region, which is characterized by similarly evolute Olcostephanus (Aguirre-Urreta & Rawson 19996). The base of the neuquensis Zone appears to lie at about the base of the Hauterivian; rare Oosterella occur either side of the boundary in both areas, while early Holcoptychites appear very close to early Spitidiscus from the lowest Hauterivian in Europe (Aguirre-Urreta & Rawson 2003) (see Fig. 8). A firmer correlation is provided in the laticosta Subzone, where the very short-ranged but widely distributed Jeannoticeras, a distinctive subgenus of Olcostephanus, occurs in the upper part. This suggests a correlation with thejeannoti Subzone. Such a correlation is supported by the discovery of very rare specimens of O. (O.) variegatus, index of the overlying variegatus Horizon in France, in the giovinei Subzone in Neuquen (Aguirre-Urreta & Rawson 2002) (Fig. 8). Spitidiscus sp. nov. (Aguirre-Urreta & Rawson 1997, fig. 7f, g) in the riccardii Zone appears almost identical to forms from the sayni Zone in France, while Crioceratites apricus (Giovine) from the schlagintweiti Zone is virtually indistinguishable from forms in the sayni-ligatus zones. The Crioceratites of the diamantense Zone have diverged from their Mediterranean counterparts in some aspects, but like the later Hauterivian-earliest Barremian Pseudothurmannia they are more tightly coiled than their predecessors. The highest zone in the Neuquen Basin has yielded only the index species, Paraspiticeras groeberi, provisionally dated as Early Barremian by Aguirre-Urreta & Rawson (1993). Calcareous nannofossil biostratigraphy Information concerning Lower Cretaceous calcareous nannofossils of the Neuquen Basin is still scarce. There have been a small number of studies with hydrocarbon exploration aims (Angelozzi 1991, 1995), and more recent papers that have focused on biostratigraphy and correlation (Mostajo et al 1995; Simeoni & Musacchio 1996; Aguirre-Urreta et al. 1999; Concheyro & Sagasti 1999; Scasso & Concheyro 1999; Concheyro et al. 2002; Concheyro & Bown 2002; Bown & Concheyro 2004). Nannofossil data compiled from several sections in the Vaca Muerta, Mulichinco and Agrio formations are presented here. They are,

LOWER CRETACEOUS BIOSTRATIGRAPHY

69

WEST MEDITERRANEAN PROVINCE

NEUQUEN BASIN

Hoedemaeker, Reboulet et a/. 2003

This paper

BIOZONE

SUB-BIOZONE (S)/ HORIZON (H)

Taveraidiscus hugii

Crioceratites diamantensis Crioceratites schlagintweiti Spitidiscus riccardii

Pleisiospit. ligatus Subsaynella sayni

Weavericeras vacaensis Olcostephanus (O.) variegatus H

O. (Jeannoticeras) jeannoti S Crioceratites loryi Crioceratites loryi S

Hoplitocrioceras gentilii

Holcoptychites neuquensis

Acanthodiscus radiatus Criosarasinella furcillata

Neocomites peregrinus Saynoceras verrucosum Busnardoites campylotoxus Thurmanniceras pertransiens

SUB-BIOZONE

Paraspiticeras groeberi

Pseudothurmannia ohmi Balearites balearis

Lyticoceras nodosoplicatum

BIOZONE

Hop. gentilii Hop. giovinei Olcostephanus (Olcostephanus) laticosta Hoi. agrioensis Hoi. neuquensis

N. (Teschenites) callidiscus S Criosarasinella furcillata S

Pseudofavrella angulatiformis

Olcostephanus (O.) nicklesi S

Neocomites sp. Chacantuceras ornatum Pseudofavrella angulatiformis O. (Viluceras) permolestus

N. peregrinus S K. pronecostatum S S. verrucosum S

Olcostephanus (Olcostephanus) atherstoni

Karakaschiceras attenuatus

K. biassalensis S Olcostephanus (O.) atherstoni

Busnardoites campylotoxus S Lissonia riveroi Neocomites wichmanni

Fig. 8. Correlation chart of the West Mediterranean and Neuquen ammonite biozones and subzones. West Mediterranean biozones after Hoedemaeker et al (2003).

Fig. 9. Calibration of Neuquen nannofossil events against the ammonite zones, and comparison with Tethyan nannofossil events calibrated against the West Mediterranean ammonite zones. Correlation across the table is based on the ammonite correlations. Tethyan nannofossil events based on Bralower et ol. (1989) and Bown et al (1998).

LOWER CRETACEOUS BIOSTRATIGRAPHY from north to south: Arroyo Cienaguitas and Arroyo Loncoche (Mendoza Province), Buta Ranquil, Pampa Tril, Cerro La Parva, Mina San Eduardo, Agua de la Mula and El Marucho (Neuquén Province) (Fig. 2). The taxonomy used here follows Bown & Concheyro (2004), where a full list of calcareous nannofossils is provided. The basal Vaca Muerta Formation is barren of nannofossils but the formation becomes nannofossiliferous up-section, yielding moderately preserved, low-diversity assemblages dominated by Watznaueria (Scasso & Concheyro 1999; Bown & Concheyro 2004). In the Pampa Tril area matrix samples from a vertebrate fossil specimen yielded Haqius circumradiatus and Micrantholiths hoschulzii (Fig. lOa); their occurrence supports the Berriasian age indicated by the ammonite Neocosmoceras sp. found at the same level. The upper part of the formation is assigned to the Lower Valanginian based on the presence of Eiffellithus primus and Eiffellithus windii, and the absence of Eiffellithus striatus (Bown & Concheyro 2004). The Mulichinco Formation is almost barren of nannofossils, with only depauperate assemblages having been reported. The Agrio Formation yields common, moderately preserved assemblages of low to moderate diversity. Nannofossil abundance increases upwards through the Pilmatué Member, which is dated as Late Valanginian and Early Hauterivian based on the presence of Eiffellithus striatus, the last occurrence of Eiffellithus windii and the occurrence of Clepsilithus maculosus (Bown & Concheyro 2004). The Avilé Member (continental or marginal marine deposits) is almost barren of nannofossils, although a poor assemblage dominated by watznauerids has been reported from the northern Mendoza Province (Angelozzi pers. Comm. 1998). Abundance and diversity increases through the overlying Agua de la Mula Member, although the topmost beds are barren of nannofossils. This member is assigned to the Upper Hauterivian based on the absence of E. striatus, the last occurrence of Cruciellipsis cuvillieri, and the presence of Speetonia colligata (Fig. lOb), Lithraphidites bollii and Nannoconus ligius (Bown & Concheyro 2004). Calcareous nannofossil bioevents Many Early Cretaceous nannofossil species have cosmopolitan distributions, and a number of global biostratigraphic events or bioevents are recognized. However, the presence of numerous,

71

more geographically restricted taxa has allowed the construction of higher resolution biozonations for both the Tethyan and Boreal realms (Thierstein 1973; Sissingh 1977; Crux 1989; Bown et al 1998; Jeremiah 2001). The nannoconids are the most abundant Tethyan nannofossil group, and they have been used in biostratigraphic schemes, particularly in Tethyan carbonate platform settings (Deres & Archéritéguy 1980; Perch-Nielsen 1985). However, the variability in their morphologies and poor preservation in limestones obscure their usefulness. Nannoconids are relatively frequent in the Neuquén Basin beds, and may have good potential for correlation (Bown & Concheyro 2004). The nannofossil bioevents determined in the Lower Cretaceous of the Neuquén Basin include the first and last occurrences (FO and LO) of Clepsilithus maculosus, Cruciellipsis cuvillieri, Eiffellithus primus, Eiffellithus windii, Eiffellithus striatus, Lithraphidites bollii and 'Nannoconus' ligius (Fig. 9). The abbreviatures NAZ indicate the local Neuquén ammonite zones, NAsZ, for Neuquén ammonite subzones, TAZ to differentiate Tethyan ammonite zones and BAZ for Boreal ammonite zones. Cruciellipsis cuvillieri bioevents The FO of Cruciellispsis cuvillieri may not be a reliable bioevent in the Neuquén Basin as it has been recorded at different levels in different sections, usually stratigraphically higher than its true earliest Berriasian evolutionary appearance (Bown et al. 1998). The FO has been recorded in Berriasian sediments at Arroyo Cienaguitas (Bown & Concheyro 2004), but in Valanginian or Hauterivian sediments elsewhere: in the wichmanni NAZ at Pampa Tril, the attenuatum NAsZ in Cerro La Parva, the ornatum NAZ at Agua de la Mula and Buta Ranquil, and the riccardii NAZ at Mina San Eduardo, where it is restricted to the Agua de la Mula Member. This rather inconsistent distribution may indicate that the species was at the edge of its ecological range, and excluded by unfavourable environments, or may reflect poorer preservation in the lower part of the sequence. At Pampa Tril, the LO of C. cuvillieri (Fig. lOc) is recorded within the riccardii NAZ (Bown & Concheyro 2004); in Agua de la Mula and Mina San Eduardo the LO coincides with the base of the schlagintweiti NAZ, and appears to be consistent with its global extinction level, corroborating the ammonite correlation with the sayni TAZ of the standard Mediterranean sequences.

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M. B. AGUIRRE-URRETA ETAL

Fig. 10. Selected calcareous nannofossils from the Vaca Muerta and Agrio formations of the Pampa Tril Section, (a) Micrantholithus hoschulzii', PTCV6, XPL. (b) Speetonia colligata, PT2, XPL. (c) Cruciellipsis cuvillieri, PTA22, XPL. (d) & (e) Eiffellithus primus PT3, PT9, XPL. (I) Eiffellithus windii, PT16, XPL. (g) & (h)Eiffellithus striatus, PTA24, PTA42, XPL. (i) Clepsilithus maculosus, PTCV6, XPL. (j) Lithraphidites bollii, PTUA 16, XPL. (k) Nannoconus ligius, PTUA26, XPL. (I) Nannoconus bucheri, PTUA27. XPL, cross-nicols, polarized light. PT indicates Pampa Tril section.

LOWER CRETACEOUS BIOSTRATIGRAPHY Eiffellithus primus bioevents The lowest occurrence of Eiffellithus primus was only recorded in the Pampa Tril area, in beds of the Vaca Muerta Formation. Its FO in Tethyan regions indicates a mid Berriasian age (Bralower et al. 1989), but in Pampa Tril the FO of E. primus (Fig. lOd, e) is in the lower wichmanni NAZ of the Lower Valanginian. The LO of E. primus occurs in the upper part of the riveroi NAZ and indicates a correlation with the Lower Valanginian campyloloxus TAZ (Bergen 1994). Eiffellithus windii bioevents The E. windii record is scattered in the studied samples, and care must be taken in distinguishing E. windii from the descendant species E. striatus. Applegate & Bergen (1989) also recognized this difficulty and chose a coccolith length of 6.4 jjim as an arbitrary morphometric cut-off between these two species. Specimens smaller than this size are included in E. windii, and we followed this definition. Eiffellithus windii (Fig. lOf) is found only in the Vaca Muerta Formation and the Pilmatué Member of the Agrio Formation, but is absent in many sections probably due to unsuitable coarse clastic marine faciès. The most complete record is at Pampa Tril, where the FO is in the wichmanni NAZ (Vaca Muerta Formation), while the LO is in the Neocomites sp. NAsZ (Pilmatué Member). Ranges elsewhere are shorter: at Buta Ranquil the FO is in the angulatiformis NasZ and the LO in the neuquensis NAsZ, while at Mina San Eduardo the species is recorded only in one sample in the Chacantuceras ornatum NasZ. The FO of E. windii and LO of E. primus are in close proximity at Pampa Tril and these bioevents have been recorded in the upper Lower Valanginian campylotoxus TAZ (Bergen 1994; Bown & Concheyro 2004). The exact level of the LO of E. windii is not well constrained in Neuquén, but at Pampa Tril it is as high as the Neocomites sp. NAsZ. The same event was recorded by Bergen (1994) in the lowermost Hauterivian radiatus TAZ. Eiffellithus striatus bioevents The FO of Eiffellithus striatus (Fig. 10g, h) has been recorded in almost all sections. At Arroyo Loncoche, Pampa Tril, Buta Ranquil and Mina San Eduardo it occurs in the angulatiformis NAsZ. At Agua de la Mula the FO is slightly higher, in the ornatum NasZ, while at

73

Cienaguitas it lies above the Neocomites sp. beds. Bergen (1994) recorded the FO of E. striatus in the Upper Valanginian trinodosum TAZ (at about the base of the furcillata Subzone of the zonation used here). At Pampa Tril the LO of E. striatus is recorded just below the Avilé Member, in the vacaensis NAZ. However, in the San Eduardo section this event occurs just above the Avilé Member (riccardii NAZ) (Concheyro et al. 2002), a level correlated with the lowermost Upper Hauterivian sayni TAZ by Aguirre-Urreta & Rawson (2001) and coincident with the established LO of E. striatus in SE France by Bergen (1994). In Agua de la Mula the LO of E. striatus is slightly higher, in the schlagintweiti NAZ. Clepsilithus maculosus bioevents Clepsilithus maculosus (Fig. lOi) ranges from the Lower Hauterivian (lamblygonium-noricum BAZs) to lowermost Barremian (lower variabilis BAZ) of the Boreal area (Bown et al 1998; Jeremiah 2001). In the Pampa Tril section its FO is recorded in the neuquensis NAZ, which is correlated with the radiatus and Crioceratites loryi TAZs (Lower Hauterivian). The LO has been determined in the vacaensis NAZ, which probably correlates with the highest part of the nodosoplicatum TAZ (Upper Lower Hauterivian) (Fig. 9). Lithraphidites bollii bioevents Lithraphidites bollii (Fig. lOj), a Hauterivian marker (Thierstein 1971; Bergen 1994), appears sporadically in some samples from Pampa Tril, Agua de la Mula and Mina San Eduardo. At Pampa Tril we only recovered fragmentary material. Its lowest occurrence is in the lower Upper Hauterivian, at the base of the riccardii NAZ; its Tethyan FO is recorded in the Lower Hauterivian loryi TAZ (Bergen 1994). At Pampa Tril an influx of L. bollii is detected in the diamantensis NAZ, and a comparable event has been recorded also in the Boreal North Sea Basin, within the gottschei BAZ (Rutledge 1995). At Agua de la Mula the distribution of L. bollii is sparse, but its FO is in the schlagintweiti NAZ and its LO in the diamantensis NAZ, in the upper part of the Agua de la Mula Member. The LO of L. bollii may not represent the extinction level, which was recorded as uppermost Hauterivian (angulicostata TAZ) by Bergen (1994).

Nannoconus ligius bioevents The problematical nannolith Nannoconus ligius (Fig. 10k) has been found in almost all the Agua

74

M. B. AGUIRRE-URRETA ET AL.

de la Mula Member sections. In Arroyo Cienaguitas and Arroyo Loncoche (Concheyro & Sagasti 1999) the FO of N. ligius is consistently above Crioceratites diamantensis up to the top of the section; while in Agua de la Mula, Mina San Eduardo and Pampa Tril the FO has been recorded within the diamantensis Zone. The LO of N. ligius occurs near the boundary between diamantensis and groeberi zones in some sections. Nannoconus ligius has been recorded previously by Applegate & Bergen (1989) from the Upper Hauterivian of the Galicia Margin. Bergen (1994) extended its range from the Upper Hauterivian (ligatus TAZ) to the Upper Aptian of SE France and the Blake-Bahama Plateau. The constant presence of N. ligius in almost all the Hauterivian Agua de la Mula Member localities of the Neuquén Basin supports the stratigraphie range originally assigned for this species by Applegate & Bergen (1989).

Biogeographic significance of nannoconids in the Neuquén Basin The presence of nannoconids, an enigmatic group of extinct nannofossils of unknown biological affinités and uncertain palaeoecology, is one of the most interesting elements of the Lower Cretaceous nannoflora of the Neuquén Basin. Their presence indicates a good marine connection with the Tethys-Caribbean seaway at this time. However, they are never as abundant in Neuquén as in contemporaneous Tethyan sequences. Nannoconids are present in all the sections that we have studied, with the exception of Cerro La Parva. They are rare and sporadic in the Vaca Muerta Formation, and not consistently present until the Pilmatué Member of the Agrio Formation (Upper Valanginian from ammonite correlations). They are absent just beneath the Avilé Member but are consistently present and often relatively common through the Agua de la Mula Member (Angelozzi 1991, 1995; Bown & Concheyro 2004). In Neuquén the nannoconid record indicates a significant lag-time between their evolutionary appearance in the Tethyan area (Tithonian) and their migration into 'extra-Tethyan' areas. But nannoconids with wide central canals, such as Nannoconus bucheri (Fig. 101), N. circularis and N. cf. N. circularis, are found relatively low in the sequence (Upper Valanginian from ammonite correlations) compared with most published ranges, which are Lower Hauterivian, Barremian and higher (Perch-Nielsen 1985; Mutterlose 1996; Bown et al 1998). These

earlier occurrences in Neuquén support a similar record by Cardin et al. (2000). However, the early occurrence of N. circularis is particularly puzzling as it is an atypically distinctive nannoconid, yet it has only been recorded previously from the Barremian to Aptian elsewhere (e.g. Deres & Archéritéguy 1980; Perch-Nielsen 1985). These observations point out to the difficulties in establishing bioevents and biozones based on this enigmatic fossil group that had a restricted biogeography.

Palynomorph biostratigraphy The oldest palynological assemblage known from the Lower Cretaceous of the Neuquén Basin was recovered from the upper part of the Vaca Muerta Formation at Mallin Quemado, c. 20km east of Las Lajas (Quattrocchio & Volkheimer 1985). The co-occurrence of the ammonite Spiticeras damesi indicates a Late Berriasian age for this assemblage. The palynoflora includes dinocysts, acritarchs and miospores. The marine assemblage includes Hystrichosphaerina neuquina (Fig. lia), a characteristic species of the Tithonian Acanthaulax downei Interval Zone, and of the Tithonian?Early Berriasian Dichadogonyaulax cumula var. curtospina Interval Zone, and Aptea notialis Range Zone (Fig. lie, d) (Quattrocchio & Sarjeant 1992; Quattrocchio et al 1996, 2002). Microforaminiferal linings are also common (Fig. llb).The terrestrial assemblages are characterized by the presence of Araucariacites australis (Fig. lie), different species of Callialasporites (Fig. I If) and Classopollis, and spores, and is comparable to equivalent assemblages recovered from the lower part of the unit and referred to the Tithonian-?Early Berriasian Microcachrydites anctarcticus Acme Zone (Quattrocchio et al 1996) and to the Late Berriasian-?Early Valanginian Zone 1 of Quatttrocchio et al (2002). The Mulichinco Formation was preliminarily studied from several different sections in the basin (Archangelsky 1977, 1980; Dellapé et al 1978; Prâmparo et al 1995; Quattrocchio et al 1999). The formation yields terrestrial and marine palynomorphs that are similar to those present in the overlying Agrio Formation, as well as marine dinocysts that are common in the Tithonian-Berriasian Vaca Muerta Formation. There are several palynological studies on the Agrio Formation. They document principally miospores and terrestrial phytoplankton (Volkheimer & Sepulveda 1976; Volkheimer

LOWER CRETACEOUS BIOSTRATIGRAPHY

75

Fig. 11. Selected palynomorphs from the Tithonian-?Early Berriasian Vaca Muerta Formation (a-d) and the Valanginian-?Early Barremian Agrio Formation (e-1). (a) Hystrichosphaerina neuquina Quattrocchio & Volkheimer emend. Quattrocchio & Sarjeant. (b) Microforaminiferal lining, (c) Dichadogonyaulax cumula var. curtospina Quattrocchio & Sarjeant. (d) Aptea notialis Quattrocchio & Sarjeant. (e) Classopollis sp. (left), and Araucariacites australis Cookson (right), (f) Callialasporites trilobatus (Balme) Dev. (g) Cyclusphaera psilata Volkheimer & Sepulveda. (h) Circulodinium distinctum (Deflandre & Cookson) Jansonius. (i) Cribroperidinium orthoceras (Eisenack) Davey. (j) Oligosphaeridium complex (White) Davey & Williams, (k) Pterospermella australiensis (Deflandre & Cookson) Eisenack. (1) Muderongia staurota Sarjeant. Scale bar is 30 (Jim.

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et al 1976; Archangelsky 1977, 1980; Volkheimer 1978, 1980; Dellapé et al 1978; Volkheimer & Quattrocchio 1981; Volkheimer & Prámparo 1984; Prámparo et al 1995; Ottone 1996; Aguirre-Urreta et al. 1996, 1999; Prámparo & Volkheimer 1996, 1999), but also marine dinoflagellates (Volkheimer & Quattrocchio 1981; Quattrocchio 1984a, b\ Quattrocchio & Volkheimer 1990; Volkheimer & Sarjeant 1993; Peralta 1994, 1996, 2000; Prámparo et al. 1995; Aguirre-Urreta et al. 1996, 1999; Prámparo & Volkheimer 1996; Mostajo & Volkheimer 1997; Peralta & Volkheimer 1997, 2000; Ottone & Pérez Loinaze 2002). The continental palynomorphs are referred to the Late Valanginian- ?Barremian Cyclusphaera psilataClassopollis sp. Palynological Assemblage (Volkheimer 1980). The palynoflora also includes Araucariacites australis, and different species of Callialasporites and Classopollis (Fig. Ile, f). Cyclusphaera psilata (Fig. llg), together with Balmeiopsis limbatus, a closely related form, display an exclusive Gondwanan distribution, whilst the rest of the miospore asemblage includes forms well known elsewhere (Aguirre-Urreta et al 1999). There are no dinocyst zonations proposed for the Agrio Formation. The only detailed palynological study was made on the Pilmatué Member of the formation at Cerro Mesa and Cerro Negro de Covunco (Peralta 1994,1996,2000; Peralta & Volkheimer 1997, 2000). The marine palynomorphs of the Agrio Formation are mostly cosmopolitan, some of them long-ranging species (e.g. Circulodinium distinctum, Cribroperidinium orthoceras, Muderongia staurota, Oligosphaeridium complex, Pterospermella australiensis: Fig. llh-1), with the exception of Muderongia brachialis, a probable local index species of the riccardii and basal schlagintweiti ammonite zones at the base of the Upper Hauterivian in the basin (Ottone & Pérez Loinaze 2002). The La Amarga Formation includes continental sediments that crop out south of Zapala. The Bañados de Caichigüe Member yields a characteristic palynological assemblage dominated by miospores and chlorococcales (Volkheimer et al 1977; Dellapé et al 1978; Prámparo & Volkheimer 2000a, b, 2002) that may be referred also to the Late Valanginian-?Barremian Cyclusphaera psilata-Classopollis sp. Palynological Assemblage. The Rayoso Group mostly comprises continental sediments, with common evaporitic horizons and red beds. A microflora recovered from the lower part of the group in the Arenisca Rincón, or Rincón Member of the Huitrín Formation (Uliana et al 1975), is characteristic of

the Late Valanginian-?Barremian Cyclusphaera psilata-Classopollis Palynological Assemblage (Vallati 1996, 2000). The palynomorph assemblage of the upper part of the group, the Ranquiles Formation, includes Cyclusphaera psilata, Balmeiopsis limbatus, Classopollis sp. and different species of Callialasporites, but also angiosperm pollen grains and Afropollis (Volkheimer & Salas 1975, 1976; Dellapé et al. 1978; Archangelsky 1980; Vallati 1995, 2000). The palynomorphs are referred to the Huitrinipollenites—Stephanocolpites Palynological Assemblage (Volkheimer et al. 1976), probably Aptian in age. The relationship between the continental palynological assemblages defined in the Neuquén Basin and the similar ones of central and southern Argentina was emphazised in several articles suggesting a great degree of floristic uniformity through the Lower Cretaceous in the region (Archangelsky et al 1981, 1984; Prámparo 1994). Discussion A detailed biostratigraphic subdivision of the Berriasian-?Early Barremian marine sequences of the Neuquén Basin has been achieved using both ammonite biozones and calcareous nannofossil bioevents. Most palynomorphs are long ranging and of only limited use in regional biostratigraphy. Correlation between the Neuquén and Mediterranean sequences and thus the accurate application of the standard stage names to the Neuquén Basin, relies on ammonite and nannofossil evidence (Figs 8 & 9). The ammonites provide firm evidence at some levels, but there are significant levels of uncertainty too. The Neuquén nannofossil species are widely distributed and many are cosmopolitan, so that the Neuquén assemblages can be directly correlated with those from the European-based stage stratotypes. Although there are only a limited number of nannofossil bioevents in the Neuquén Basin, they provide a robust independent test for the correlations indicated by ammonites. In general, the ammonite correlations are supported by nannofossil evidence. This is especially true from the noduliferum to the atherstoni NAZs (Early Berriasian-Upper Valanginian). But the occurrence of wide-canalled nannoconids in the angulatiformis NAsZ could suggest a slightly younger, Early Hauterivian, age for this zone than is indicated by the ammonite correlations. However, Bown & Concheyro (2004) do not place great confidence in the calibration of these nannoconid events, while the

LOWER CRETACEOUS BIOSTRATIGRAPHY first occurrence of Eifellithus striatus at the base of the angulatiformis NasZ supports the ammonite correlation. Although the base of Tethyan nannofossil zone CC4B is higher in the Neuquén Basin according to the ammonite correlations, there is a remarkably good agreement in the mid Hauterivian, particularly the LO of Cruciellipsis cuvillieri at the base of the schlagintweiti NAZ, which corroborates the ammonite correlation with the sayni TAZ of the European sequences. Finally, the LO of Nannoconus ligius in the basal part of the Paraspiticeras groeberi NAZ suggests a Late Hauterivian age for at least part of this ammonite zone. As noted above, Aguirre-Urreta & Rawson (1993) provisionally dated the zone as Early Barremian, but part or all could be latest Hauterivian as that is when the genus Paraspiticeras first appeared. Our team research was supported by a grant from the British Council/Fundación Antorchas. M.B. AguirreUrreta acknowledges financial support from ANPCYT 14143/03 and UBACyT x 084, and P.P. Rawson acknowledges the award of a 2003 Leverhulme Emeritus Fellowship from The Leverhulme Trust. E.G. Ottone is grateful to G.D. Holfeltz for laboratory assistance and helping with the photography. G.A. Concheyro is grateful to G. Villa for fruitful discussions about taxonomy during the author's scholarship at the University of Parma, Italy.

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Cretaceous nannofossil marker-species. Journal of Nannoplankton Research, 18, 53-59. SCASSO, R. & CONCHEYRO, A. 1999 Nanofósiles calcáreos, duración y origen de ciclos calizamarga (Jurásico tardío de la cuenca Neuquina). Revista de la Asociación Geológica Argentina, 54, 290-297. SIMEONI, M. & MUSACCHIO, E.A. 1996.Cretaceous calcareous microfossils from Southern South America: paleobiogeographic relationships. Zentralblatt für Géologie und Palaontologie, I, (11/12), 1403-1420. SIMONESCU, I. 1899. Note sur quelques Ammonites du Néocomien français. Travaux Laboratorie Géologique, Faculté des Sciences, Université du Grenoble, 5, 1-17. SISSINGH, W. 1977. Biostratigraphy of Cretaceous calcareous nannoplankton. Géologie en Mijnbouw, 56, 37-65. THIERSTEIN, H. 1971. Tentative Lower Cretaceous calcareous nanoplankton zonation. Eclogae Geolgicae Helvetiae, 64, 459-488. THIERSTEIN, H. 1973. Lower Cretaceous Calcareous nannoplankton biostratigraphy. Abhandlungen der Géologie Bundesanstalt, 29, 1-52. ULIANA, M.A., DELLAPÉ, D.A. & PANDO, G.A. 1975. Estratigrafía, distribución y génesis de los depósitos rayosianos. 77° Congreso Iberoamericano de Geología, Actas, 1, 151-170. VALLATI, P. 1995. Una microflora con Afropollis (polen de angiosperma) en el Cretácico Inferior de la Cuenca Neuquina. VIo Congreso Argentino de Paleontología y Bioestratigrafía, Actas, I, 277-290. VALLATI, P. 1996. Palinología de la 'Arenisca Rincón' (Grupo Huitrín-Rayoso), Cretácico Inferior de la Cuenca Neuquina. X777° Congreso Geológico Argentino & IIIo Congreso de Exploración de Hidrocarburos, Actas, 5, 77-92. VALLATI, P. 2000. Recurrencia litológica vs. cambios evolutivos: un caso en el Cretácico de la Cuenca Neuquina. llth Simposio Argentino de Paleobotánica y Palinología, Resúmenes, 1, 110. VOLKHEIMER, W. 1978. Microfloras fósiles. In: Geología y Recursos Naturales de la provincia del Neuquén. 7th Congreso Geológico Argentino, Relatorio. Asociación Geológica Argentina, Buenos Aires, 193-207. VOLKHEIMER, W. 1980. Microfloras del Jurásico Superior y Cretácico Inferior de América Latina. 77° Congreso Argentino de Paleontología y Bioestratigrafía & 7° Congreso Latinoamericano de Paleontología, Actas, 5, 121-136. VOLKHEIMER, W. & PRÁMPARO, M.B. 1984. Datos palinológicos del Cretácico Inferior en el borde austral de la Cuenca Neuquina, localidad Estancia Santa Elena, Argentina. Parte I: especies terrestres. In: 3rd Congreso Latinoamericano de Paleontología, Memoria. UNAM, México, 269-279. VOLKHEIMER, W. & QUATTROCCHIO, M.E. 1981. Distribución estratigráfica de los palinomorfos jurásicos y cretácicos en la faja andina y áreas adyacentes de América del Sur austral con especial consideración de la Cuenca Neuquina. In:

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Sedimentology of the tide-dominated Jurassic Lajas Formation, Neuquen Basin, Argentina DUNCAN McILROY1, STEPHEN FLINT2, JOHN A. HOWELL3 & NICK TIMMS2 1 Department of Earth Sciences, Memorial University of Newfoundland, St John's, Newfoundland, Canada A1B 3X5 (e-mail: [email protected]) ^Department of Earth Sciences, University of Liverpool, 4 Brownlow Street, Liverpool L69 3GP, UK 3 Geological Institute, University of Bergen, Allegaten 41, N-5007 Bergen, Norway Abstract: Tidal depositional systems are often interpreted as lowstand/transgressive estuarine deposits within sequences that are either wave or river dominated during highstand times. The Middle Jurassic Lajas Formation of the Neuquen Basin, Argentina, comprises 600 m of well-exposed tide-dominated facies deposited within four unconformitybounded sequences, spanning approximately 4.5 Ma. Facies associations include tidedominated deltas, sandy-heterolithic tidal channel fills and extensive progradational tidal-flat successions, which are locally cut by heterolithic tidal channel fills. Despite the narrow bathymetric depositional range and the complex facies variability, flooding surfaces can be defined and mapped along a 48 km-long outcrop belt. These flooding surfaces allow definition of three distinct types of parasequence that exhibit coarsening-upwards, finingupwards and coarsening- to fining-upwards motifs. Sequence boundaries are marked by widespread, but shallow, incision, and the juxtaposition of stacked fluvial/tidal channel fills on a variety of subtidal and intertidal facies. Unconventional grain-size changes at sequence boundaries can occur where basinward facies shifts are marked by juxtaposition of heterolithic-argillaceous intertidal/supratidal mudflat deposits on subtidal sandflat facies. The maintenance of macrotidal conditions through complete base-level cycles is interpreted as being due to the structural topography inherited from rifting, causing the whole sub-basin to behave as a structurally controlled embayment.

The majority of modern and ancient tidal depositional systems described in the literature occur in estuarine settings. Estuaries are primarily transgressive features (Dalrympleer al 1992) that represent the flooding of valleys cut during periods of low sea level. The morphology of estuaries amplifies tidal currents and provides shelter from wave reworking (Allen 1991; Dalrymple et al. 1992). The dominantly transgressive nature of estuary fills precludes large-scale progradation of tidedominated facies. Once valley systems are completely filled, the sedimentary system typically reverts to wave- or river-dominated open coastlines due to a lack of tidal amplification. This spatial and temporal restriction of significant volumes of tidal facies to estuaries and times of early base-level rise is an important predictive element of sequence stratigraphic models, and has proven successful in explaining many ancient mixed-tidal and open-shelf deposits (Posamentier & Vail 1988; Shanley et al. 1992). However, this model does not adequately

explain: (a) present-day progradational, nonestuarine tide-dominated systems, such as the Ganges-Brahmaputra and Fly River deltas; and (b) extensive ancient progradational tidedominated successions up to hundreds of metres thick, such as the Jurassic Tilje and He formations, offshore mid-Norway (Martinius et al. 2000; Mcllroy 2004), and equivalent outcrops in Jameson Land, Greenland (Dam & Surlyk 1995). Similar successions have been documented in small fault-bounded basins in Scotland (Mellere & Steel 1996) and in lowstand deposits from the Cretaceous of North America (Willis et al. 1999; Bhattacharaya & Willis 2001; Willis & Gabel 2001). In this paper we describe and interpret the sedimentology of a 600 m-thick, largely aggradational, completely tide-dominated succession of Middle Jurassic age from the Neuquen Basin, Argentina and propose a sequence stratigraphic model for meso-/macrotidal successions, which incorporates tidal deltaic sedimentation in both incised valley and highstand deltaic settings.

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 83-107. 0305-8719/05/$15.00 © The Geological Society of London 2005.

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Geological setting The Neuquen Basin of western Argentina (Fig. 1) was formed in a back-arc extensional setting during the early stages of the Andean Orogen (Digregorio & Uliana 1980). The basin inherited its distinctive triangular geometry from Palaeozoic terrane sutures (Uliana & Riddle 1987). The earliest (Triassic) phase of sedimentation included deposition of epiclastic volcanic, and continental red-bed facies in half-grabens. These fault-bounded depocentres became progressively interconnected prior to a widespread marine flooding event in the Pliensbachian Stage. The basin then remained dominantly marine until structural inversion in the Callovian Stage (Legaretta & Uliana 1991). LowerMiddle Jurassic deposits of the Cuyo Group (Fig. 2) in the Sierra del Chacaico (Fig. 1) comprise approximately 2000 m of marine and

continental strata. The lower Cuyo Group is characterized by pelagic shales and turbidites of the Los Molles Formation that, in the central and northern parts of the basin, persist through to the Callovian Stage (Gulisano & Gutierrez Pleimling 1994). The Los Molles Formation is diachronously overlain by shallow-marine deposits of the Cura Niyeu-Lajas Formation shelf system (Fig. 2), which prograded from the southern and eastern margins of the basin (Gulisano & Gutierrez Pleimling 1994). Previous sedimentological studies of the Lajas Formation identified its tide-dominated character and used estuarine facies models to explain the abundance of tidal sediments (Zavala 1996), but no formal facies associations were proposed. A preliminary account of the ichnology has also been presented by Poire & del Valle (1992). Initial basin-scale sequence stratigraphic

Fig. 1. Maps showing the location of the study area and the outline of the Neuquen Basin. The main studied areas are highlighted by three boxes, of which Box A is detailed in Figure 3. Field area A is located within a pair of faults that are inferred to have enclosed a triangular embayment on the southern margin of the Neuquen Basin in the Jurassic. Fault lineaments after Vergani et al. (1995).

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B, Fig. 1) to 180m at Estancia Charahuilla (excluding the underlying Cura Niyeu Formation). Superposition of Cuyo Group outcrops on the fault map of Vergani et al. (1995) shows that all tide-dominated facies belts lie within a N-S-orientated, fault-defined embayment (Fig. 1). The western margin of the embayment is coincident with the Sierra de Catan Lil and the eastern margin is to the east of the outcrop belt in the Sierra del Chacaico (Fig. 1). In this structural embayment (area A, Fig. 1) up to 600 m of tidal sediments accumulated over a period of around 4.5 Ma. At the north end of the embayment, the Lajas Formation is typified by dm-scale dunes and some storm deposits (area B; Dean 1987 and this study). Lajas facies, that crop out to the north (seaward) of the embayment (area C, Fig. 1), are wave/storm-dominated rather than tidal in nature. However, palynological data in Quattrochio & Volkheimer (1990) suggest that the Lajas Formation in this area may be considerably younger than in the main study area. Fig. 2. Jurassic stratigraphy of the Neuquen Basin showing the position of the Cura Niyeu/Lajas/Challaco depositional system overlying the basin and slope deposits of the Los Molles Formation.

interpretations (Gulisano & Gutierrez Pleimling 1994; Zavala 1996) interpreted channel-prone parts of the Lajas Formation as estuarine incisedvalley fills overlying sequence boundaries. A preliminary account of Lajas Formation facies and parasequence architecture has been presented by Mcllroy et al (1999). The Lajas Formation is diachronously overlain by the nonmarine Challaco Formation (Gulisano & Gutierrez Pleimling 1994). The Challaco Formation is comprised of coarse-grained fluvial-channel deposits and mudstone-rich floodplain deposits. This study concentrates on the description and interpretation of a high-resolution sedimentological dataset collected from a superbly exposed 48 km-long cliff line (area A, Fig. 1). Sections were measured at 1 km spacing, with shorter 'infill' logs at important localities (Fig. 3). Correlation was performed by tracing key stratigraphic surfaces between logged sections, and was aided by the integration of data from photomontages and oblique-aerial photographs. Regional setting and palaeogeography The thickness of the Lajas Formation decreases southward from 800 m at Arroyo Carreri (area

Sedimentology Ten facies associations have been defined based on sedimentological and ichnological criteria, and are summarized in Figure 4. Bioturbation was quantified using the bioturbation index (BI) of Taylor & Goldring (1993); following Reineck (1963, 1967). The facies associations are related to palaeoenvironmental interpretations based on the vertical succession of facies and tracing lateral geometries over several kilometres. The following discussion of facies associations is organized from distal (offshore marine) to proximal (tidally influenced fluvial). The sedimentological descriptions below are supplemented by reference to a small-scale study of facies architecture (Fig. 5). Facies association I (FA1): offshore and tide-dominated delta

shelf

FA la: shelf mudstones and prodelta heterolithic strata Description. This facies association consists of mudstones containing ammonites, belemnites, marine reptiles, fragments of high-spired gastropods and thin-shelled bivalves of ? Modiolus. There is no trace of internal lamination due to intense bioturbation (BI 5-6), which also precludes identification of discrete trace fossils. Sandstone beds up to 4 m in thickness contain partial Bouma sequences (Fig. 4) but comprise

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Fig. 3. Geology of the Sierra de Chacaico (field area A in Fig. 1) showing the location of logged sections and the location of the detailed study area that comprises Figure 5.

a comparatively small part of the facies (around 20%) and are typically 10-30 cm in thickness. Interpretation. These mud-rich deposits were deposited on an offshore shelf, where the rate of sedimentation was low in proportion to infaunal biological activity. The lateral and vertical relationships with tidal delta deposits (FA Ib) implies a subtidal prodeltaic origin. The sandstone interbeds are interpreted as shelfal, deltafront turbidites. In the lower part of the Cura Niyeu Formation there is a gradual upwards transition from offshore mudstones of the underlying Los Molles Formation into this facies association. Within the main Lajas Formation, FA la

is poorly represented due to erosion during delta-plain progradation (e.g. Perro Loco parasequence, logs 7-17; Fig. 5) (cf. Mcllroy 2004). FA Ib: tide-dominated delta front Description. This facies association occurs as 10-15m-thick units in the basal stratigraphic section (upper Cura Niyeu Formation) and as 4-5 m-thick units in the middle part of the Lajas Formation. Pro-delta mudstones of FA la coarsen upwards through heavily bioturbated (BI 5-6) siltstone and sandy siltstone (containing Ophiomorpha, Thalassinoides, Asterosoma, Rosselia, Schaubcylindrichnus, Palaeophycus

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Fig. 4. Summary diagram showing the range of shallow-marine tidal deltaic to fluvial facies associations described in this study. The facies are diagrammatically represented in a broadly progradational succession that is not intended to represent any real stratigraphic succession.

and Parahaentzschelinia) into 1-10 cm-thick beds of fine-grained, planar-bedded sandstone interbedded with bioturbated siltstone (Fig. 4). This heterolithic facies coarsens upwards into

approximately 50 cm-thick stacked beds of heavily bioturbated (BI4-5) cross-bedded sandstones or larger cross-sets up to 8 m thick with sigmoidal geometries (Fig. 6a). Superimposed

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Fig. 5. Detailed facies architectural panel showing facies variation within three parasequences from which estimates of facies dimensions and interrelationships can be gained.

TIDAL SEQUENCE STRATIGRAPHY

bedsets are smaller ripple/dune cross-bedded showing bidirectional palaeocurrents orthogonal to the larger cross-sets. These cross-bedded facies are moderately bioturbated (BI 2-3) and commonly show paired drapes defined by concentrations of fine participate carbonaceous debris. These thinly bedded-cross-bedded facies can be seem to be associated with larger scale structures with a clinoform geometry, dipping northward at 6° (see Fig. 13 later in this paper). Medium- to coarse-grained sandstones, with 10°-15° dipping large-scale crossbeds or massive irregular-based sands typically erode the tops of complete upwards-coarsening successions (Fig. 6b). Interpretation. In thick examples of this facies association, characteristic of the base Lajas/Cura Niyeu formations, the northward-dipping clinoforms are interpreted as delta-front foresets of prograding tide-dominated deltas. The delta front is characterized by low-relief barforms, which become more pronounced upwards. In the thinner examples of this facies association, within the Lajas Formation proper, cross-sets are generally orientated orthogonal to the mean channel axis, probably representing laterally

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migrating tidal distributary mouthbars (Fig. 6a). The depth of erosion at the contact between mouthbars and their feeder channels (Fig 6b) controls the preservation potential of the tidal mouthbars. This interaction between bar and feeder channel is analogous to the bay head diastem of Nichol et al. (1994) and is a normal autocyclic process that acts during progradation of the delta system (e.g. lowest part of the Perro Loco parasequence, logs 1-6; Fig. 5). FA Ic: bay-fills Description. Successions of heavily bioturbated/burrow mottled (BI 5-6) mudstones and sandy mudstones up to 15m thick are relatively common in the Lajas Formation and crop out continuously over 1-4 km. This facies association also contains shell beds up to 3 m thick (Fig. 7), composed of several species of thickwalled in situ oysters up to 10 cm long with shells up to 4 cm thick, exhibiting Gastrochaenolites isp. borings. In places, these heavily bioturbated mudstones coarsen and thicken upwards into medium-grained muddy sandstones with wave and current ripples. The lateral margins of this facies association are typically marked

Fig. 6. Facies of FA 1 (tide-dominated deltas): (A) steep deltaic bedsets showing little erosion of the topsets (geologist for scale bottom right); (B) channel-form with a homogeneous fill cutting into the top of an upwards-coarsening deltaic succession such as seen in (A) above (increments on measuring staff 10 cm).

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D. McILROYCTAL. Facies association 2 (FA2): tidal channels

Fig. 7. Mud-rich succession in the middle portion of the Lajas Formation containing a thick autochthonous shell bed (figure for scale) interpreted as bayfill deposits (FA le).

by erosion by overlying tidal channels. Where a non-eroded margin is exposed (e.g. the Komplott parasequence, N of log 16; Fig. 5) it grades laterally into a tidal-flat facies association 3a/b.

Interpretation. The fine grain size and intensity of bioturbation suggest deposition largely from suspension away from major axes of tidal-delta progradation. The presence of thick beds of large oysters is suggestive of good advective mixing between sea water and freshwater. Such well-mixed environments, capable of supporting marine taxa for protracted periods, can be found in large restricted, 'leaky' embayments/lagoons (sensu Kjerve & Magill 1989) or large bays. The latter environment is preferred because of a lack of evidence for barrier bars or wash-over fans in the Lajas, which are required for evocation of a lagoonal setting. The shallow marginal marine interpretation of this facies is further supported by the fact that this facies can be traced laterally into tidal-flat facies FA 3a/b.

FA 2a: meandering sandy tidal channels Description. This facies association is characterized by erosively based, medium- to coarsegrained sandstones, comprising 2-4 m-thick, inclined cross-sets (10°-30°) with superimposed trough cross-bedding (Fig. 8a). Palaeocurrents from the trough cross-beds are approximately perpendicular to the dip of the parent crossbeds. Toesets of the low-angle major cross-beds commonly contain current ripples, mostly exhibiting northward (ebb-oriented) palaeocurrents. Bases of units are moderately bioturbated (BI 2-3) by Scolicia isp. andAsteriacites, and contain thin interbeds of mudstone. The upper parts of the cross-sets do not have significant clay drapes and are much less bioturbated (BI 1) with scattered clusters of Dactyloidites isp. This sandrich cross-bedded facies may form multistorey units up to 10m thick and 600 m wide. The facies association typically erosively overlies and erodes into FAs 1 and 3 (e.g. Komplott parasequence, log 7; Fig. 5). Silicified wood fragments with Teredolites isp. borings are also occasionally found (Morgans-Bell & Mcllroy 2005). A variant of this facies association comprises medium- to coarse-grained, well-sorted, trough/planar cross-bedded sandstone, with either single or paired drapes, which are defined by concentrations of organic debris that resemble 'coffee grounds' (cf. Shanley et al 1992). Trough cross-sets are rarely thicker than 30 cm, whereas the planar cross-sets are commonly 1 m thick. The drapes commonly show rhythmic bundling, defined by densely spaced draped toesets separated by less distinct, undraped foresets (Fig. 8b). Cross-sets with single drapes defined by comminuted organic debris are typically trough cross-bedded, grading into larger planar cross-beds with paired organic drapes (Fig. 8c), wavy-flaser bedded foresets and reactivation surfaces. Palaeocurrents are typically unimodal (ebbdominated) in the trough cross-bedded facies, but may be bimodal in the planar cross-bedded facies. Bioturbation is generally sparse (BI 12); rare Planolites burrows are concentrated in toeset drapes and occasional Dactyloidites isp. are found in the planar cross-bedded facies. Interpretation. The low-angle inclined surfaces are highly oblique to palaeocurrent indicators measured from superimposed trough-crossbedding, which suggests that this facies association is dominated by lateral accretion deposits formed on migrating point bars within

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Fig. 8. Facies of FA 2 (tidal channels): (A) multistorey tidal channel complex composed of meandering tidal channel faciès (FA 2a) (figure for scale bottom centre); (B) trough cross-bedding with tidal bundling composed of fine organic debris on foresets deposited during slackwater in the neap part of the tidal cycle (FA 2a); (C) detail of tidal mud-drapes showing couplets that suggest a subtidal depositional environment; (D) heterolithic tidal-channel fill (FA 2b) showing steeply inclined lateral accretion deposits comprised of thin ripple cross-laminated sandstone and organic-rich mudstone.

meandering channels. The presence of mud couplets, bundling, reactivation surfaces and bimodal palaeocurrents indicate a strong tidal influence. The trough and planar cross-bedded faciès with double drapes and bimodal palaeocurrents are interpreted as the subtidal portion of tidal channels (cf. Visser 1980; Smith 19886; de Boer et al 1989; Fenies et al 1999). These sandstones are interpreted as compound barforms, preserved in the inner bank of meandering tidal channels in the manner described by Yang & Nio (1989). The trough crossbedded faciès with single drapes and pronounced tidal bundling is interpreted as intertidal. The low intensities of bioturbation recorded in these faciès probably relate to fluctuating salinity and high current velocities experienced in fluvially attached tidal channels through much of the tidal cycle, which discouraged suspensionfeeding infauna. However, the presence of

echinoderms, as indicated by the trace fossils Scolicia isp. and Asteriacites, demonstrates that a normal marine infauna was present at least ephemerally and that porewaters were dominantly marine in character. This is in line with recent observations on the salinity tolerance of the echinodermata (Mângano 1999; Mângano & Buatois 2004). The grain size of tidal-channel sandstones is typically slightly coarser than in delta-front and tidal-flat faciès. This storage of coarser grained sediment in meander-belt systems is documented from the modern Fly River Delta, which has an extensive tidal plain up to 15 km wide, where most coarse-grained sediment is trapped in meander-belt deposits (Loffler 1977; Harris et al 1993; Baker et al 1995). Similarly, the Ord River in Western Australia has a meander belt 10 km wide, cutting a tidal plain covered by tidal flats (Coleman & Wright 1978).

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FA 2b: meandering heterolithic tidal channels Description. This facies association consists of interbedded fine-grained current-rippled sandstones up to 5 cm thick and mudstones up to 3 cm thick, forming bedsets up to 3 m thick that dip at 10°-15° (Fig. 8d). Palaeocurrents from ripples in the sand-rich beds are typically oriented perpendicular to the inclined crossbeds and are bimodal. The cross-sets are typically moderately bioturbated (BI1 -2), with some Planolites and unidentifiable bioturbation. The main difference from FA 2a is that the sandstone portion is finer grained and the mudstone layers extend to the tops of the low-angle dipping surfaces. This facies is commonly associated with thick tidal-flat successions (described below). Interpretation. The angle of inclination of the accretionary beds and heterolithic nature of this facies association is comparable to that described from both modern and ancient inclined heterolithic strata (IMS) sensu Thomas et al. (1987), known to be deposited in tidal point bars. The sparse bioturbation suggests deposition in stressed, possibly intertidal, palaeoenvironments. Examples of mud-rich tidal point bars from Ecuador (Smith 1988a) show a similar absence of bioturbation in the majority of inclined bedsets, except at their tops. Sand-grade sediment was probably sourced from ebb-tidal draining of adjacent tidal flats and/or flood-tidal processes, rather than from an attached fluvial source (cf. Bridges & Leeder 1976; de Mowbray & Visser 1984; Smith 1988£; Shanley et al 1992). FA2c: bioturbated, abandoned channel fills Description. Intensely bioturbated facies are occasionally found in association with the largely unbioturbated tidal-channel facies associations. Examples are typically 2-3 m thick, rarely greater than 150 m wide and commonly eroded by sandy tidal channels. Little primary sedimentary lamination is preserved (i.e. BI 4-6) and the mud-dominated facies (Fig. 9a) is generally rich in detrital organic matter and includes the trace fossils Thalassinoides, Teichichnus, Ophiomorpha, Rosselia, Palaeophycus and Asterosoma. The sand-dominated equivalent is characterized by a monospecific ichnofauna of Ophiomorpha isp., with burrow walls lined with organic detritus and a BI of 4-5. Interpretation. The close spatial relationship with the sandy tidal channels implies a genetic linkage. Lithofacies trends are similar to those

described by Sha & de Boer (1991) from abandonment of - previouslyfluviallyattached tidal-channel deposits in which the fluvial input has been cut off. The resultant lack of freshwater input allows a normal marine biota to colonize the channel, generating ichnofabrics of marine affinity, which means that these facies can be mistakenly interpreted as marine flooding surfaces (see below). Similarly, in ebb-dominated parts of the modern Fly River Delta, it is only on the largely abandoned/flood-dominated NE edge of the delta that significant marine biota is found in the tidal channels (Alongi 1991). This is in contrast to FA2b, in which the channels may completely drain to low tide, thereby precluding establishment of a permanent marine infauna. Facies association 3 (FA3): tidal flats FA 3 a: intertidal muddy-mixed flats Description. This heterolithic association consists of interbedded mudstones and fine- to medium-grained, wave- and current-rippled sandstones (Figs 4 & 9b). Single clay drapes on ripple cross-bedding are common and desiccation cracks are also recorded. There is a complete continuum between mud- and sand-dominated heterolithic facies, with the thickest sandstone beds commonly occurring in facies with the lowest mud to sand ratio. This facies association is generally rich in detrital plant matter and may be cut by impersistant, poorly sorted, crossbedded sandstones. Bioturbation is typically of low diversity and intensity (BI 1-2), although moderate bioturbation (BI 3) may occur in the sand-dominated end member. No paired clay drapes were found, although single drapes forming tidal bundles are locally common in the sand-rich end member. The muddy end member of this facies association is comparatively rare, areally restricted and is usually eroded by tidal-channel facies (Komplott Parasequence, log 10, Fig. 5). Interpretation. Abundant single clay drapes in cross-bedded sandstones represent deposition at high-tide slack water (subtidal settings are characterized by double drapes, representing the slack-water phase of both the high and low tides). The extensive, almost flat-bedded sheet-like geometry (traceable over several kilometres), lack of bioturbation and single drapes evince a muddy-mixed intertidal-flat environment. The erosive-based, tangentially based cross-bedded units are interpreted as the deposits of meandering tidal creeks that drained the tidal flat during the ebb portion of the tidal

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cycle (e.g. Bridges & Leeder 1976). Terwindt (1988) cites the preservation of thick intertidalflat successions with abundant intertidal drainage creeks as diagnostic evidence of a meso- to macrotidal regime. The extensive tidal-flat facies in the macrotidal Kyeonggi Bay, Inchon, Korea (Frey et al 1989; Kim et al 1999) may also form a good modern analogue for the broad tracts of tidal-flat facies in the Lajas. The Korean tidal flats, however, are in an estuarine setting and pass distally into subtidal channels. In contrast, the Lajas Formation tidal flats pass offshore through subtidal flats (FA 3c) into shelf mudstones (Facies 1A) and are not subject to the rapid salinity fluctuations characteristic of estuaries. FA 3b: intertidal sandflats Description. Sand-rich facies with 10-30 cmthick cross-sets, showing well-developed tidal

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bundles defined by single clay/organic-rich drapes (Fig. 9c) represent approximately 2030% of the Lajas Formation (Fig. 4). The sandstones are fine to medium grained, and show variable proportions of wave- and tideinfluenced sedimentary structures (Fig. 9d). Bioturbation is uncommon but, where present, is normally intense (BI 4-5) and of low diversity (particularly Planolites isp. and Dactyloidites isp.). These deposits typically form laterally extensive sheets that are extremely heterogeneous in nature. In cases where this facies lies adjacent to tidal-channel facies (e.g. several localities in logs 1-5 in both the Perro Loco and Komplott parasequences) they tend to be slightly coarser (medium-coarse) grained, show little evidence of wave-generated structures and are dominated by trough cross-beds with tidal bundling and herringbone cross-stratification. Away from tidal-channel-dominated areas, wave ripples are a common component,

Fig. 9. (A) Highly bioturbated channel-fill deposit with branching Thalassinoides burrows at the base of the flaserbedded late-stage channel fill. Interpreted as an abandoned tidal-channel deposit (FA 2c) (camera lens 5.5 cm in diameter). (B)-(D) represent tidal-flat facies from the Lajas Formation (FA 3): (B) shows a mud-rich deposit with thin wave and current ripple cross-laminated sandstones that is interpreted a mudflat deposit (FA 3a) (camera lens 5.5 cm in diameter); (C) sand-rich thinly bedded facies with small-scale trough cross-bedding and abundant wave-ripples but with only sparse draping of foresets. Interpreted as a strongly wave influenced tidal flat deposit (FA 3b) (increments on measuring staff are 10 cm); (D) sand-rich, thinly bedded heterogeneous facies with abundant drapes on ripple crosslamination and trough cross-bedding. Sedimentary structures produced by oscillation ripples are rare (increments on measuring staff are 10 cm).

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although tidal bundling and clay/organic matter draped foresets can still be recognized. Interpretation. This facies association is interpreted to have been deposited in intertidal environments close to tidal channels, passing laterally into more wave-influenced settings (Fig. 10) with more stable, normal marine salinities due to an absence of local freshwater input. Sediment was probably largely transported to the coastline by tidal channels and subsequently redistributed by coast-parallel tide and wave processes, rather than being supplied from the shelf as in many modern estuarine systems. The lowered ichnodiversity is probably a function of significant temperature stress acting during the extensive neap-tidal periods in such a warm climate with a large tidal range. FA 3c: subtidal flats Description. This facies association consists of fine-grained sandstones with thin clay-draped wave/current ripples. Some horizons contain bedding surfaces that dip at up to 5°, characterized by simple flaser bedding, with mudstone

drapes preserved only in ripple troughs. Occasionally, beds show rhythmic alternations between mud- and sand-rich deposition. Bioturbation (BI 4-5) is characterized by Teichichnus and Rosselia. This facies is found in erosive contact with subtidal cross-bedded channel fills (Perro Loco Parasequence logs 13 and 14; Fig. 5) and in gradational contact with programing coastal flats without extensive channel deposits. Interpretation. These bioturbated heterolithic strata are interpreted as subtidal flats with lowamplitude, long-wavelength shoals (positive relief non-migrating features) that overlie and pass distally into bioturbated offshore mudstones of FA la. Similar subtidal facies may also be found associated with the subtidal fill of large tidal channels, in which sediment supply is inferred to have come from the channel margin rather than from a tide-dominated coastline. Facies association 4 (FA 4): fluvial channels andfloodplain Description. Examples of fluvial deposits within the Lajas Formation are volumetrically

Fig. 10. Conceptual facies model based on the facies distributions in Figure 5 showing the relationship between tidalflat and tidal-channel environments. Note that tidal flats are commonly developed on top of tidal-channel meander-belt facies.

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rare and comprise coarse- to very-coarse-grained sandstones less than 1 m thick and overlain by stacked tidal channels. The basal surfaces of the units are erosional and they trace laterally to poorly developed palaeosols. Overlying the Lajas Formation is the fluvially dominated Challaco Formation, the base of which is taken at the first occurrence of stacked coarse-grained largescale cross-bedded units or thick palaeosols. The base of the Challaco Formation is highly diachronous (Fig. 2). In the south of the study area the Challaco was deposited contemporaneously with tidal facies of the uppermost Lajas Formation further to the north (see the Sequence stratigraphy section later in this paper). The transition between the Lajas Formation and the Challaco Formation in the north of the Sierra de Chacaico is marked by up to 20 m of amalgamated, channelized, trough cross-bedded immature sandstones and conglomerates (Figs 4 & 11 a). The basal cross-bedded units commonly contain silicified logs of the araucarian woodgenus Araucarioxylon sp. (Morgans-Bell & Mcllroy 2005 (Fig. lib), with rare bivalve borings. Overlying facies are mudstone-

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dominated with colours ranging from black, green, blue and purple to red. These coloured mudstones commonly contain rootlet beds (Fig. lie) and current rippled 3-10 cmthick planar-bedded sandstones, both of which may be cut by isolated trough cross-bedded sandstone/conglomeratic units up to 8 m thick and 200 m wide. Interpretation. The presence of large-scale erosional features filled by cross-bedded texturally and mineralogically immature sandstone suggests that the main Challaco Formation sandstones were deposited in fluvial channel systems. Evidence for limited marine influence is provided by bivalve borings in the fossil wood found within these channels. The up-dip passage from meandering tidal channels into fluvial channels is interpreted to be gradual, owing to the large tidal reach expected in such a system, although such a transition cannot be traced out in the field due to insufficient exposure. The coloured mudstones with rootlets are interpreted as floodplain palaeosols and the rippled sheet sandstones as minor crevasse-splay deposits.

Fig. 11. Sedimentary facies comprising facies association 4 (FA 4): (A) the base of the Challaco Formation comprising steep cliffs of tidally influenced coarse-grained fluvial channels (cliffs are approximately 10 m high); (B) contorted piece of silicified wood from the Lajas/Challaco transition showing twisting characteristic of driftwood; (C) rootlet bed from the Lajas Formation forming an incipient soil horizon with complete disruption of the primary tidalchannel/tidal-flat sandstone fabric.

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Non-marine facies within the main part of the Lajas Formation, although rare, are important stratigraphic markers. They are interpreted to represent periods of marked progradation of the fluvial systems into the tide-dominated area of the Lajas embayment and their significance is discussed further below.

Lajas Formation holistic facies model The Lajas Formation contains tidal mouth bars similar to those found in classic tide-dominated estuaries (cf. Dalrymple et al. 1992) and deltas (e.g. Willis et al 1999; Willis & White 2000), but differs from most facies models for macrotidal settings in the extensive preservation of tidalflat deposits and the progradational character of the system. In estuaries, the laterally confined nature of systems generally results in the complete cannibalization of tidal-flats, because the width of tidal-channel meander belts is commonly equivalent to the valley width. Much of the Lajas coastline is inferred to have been dominated by prograding tidal flats like those of the German Bight (Reineck & Gerdes 1996; Kim et al. 1999) (Fig. 12). In settings such as these, tidal-channel meander belts are rarely wide enough to destroy all of the intervening tidalcoastline deposits (mainly tidal-flat and bayfill facies). The dimensions of these meander belt sandstones are thus dependent on the rate of channel migration and its evolution through time, particularly with respect to the rate of accommodation generation (Mcllroy etal 1999).

Lajas Formation tidal-flat deposits are similar to coastally attached, fining-upwards tidal barforms recognized by Klein (1971) as prograding tidal flats. An idealized progradational tidal-flat succession has at its base offshore marine mudstones (FA la) overlain by heterolithic, subtidal flats (FA 3c). These coarsen upwards (i.e. landward) into sandy intertidal flats that in turn fine upwards into heterolithic mixed-muddy tidal flats. The idealized progradational succession is therefore heterolithic at its base and top, with sand deposition on tidal flats in the lowermiddle part of the intertidal zone. Some sections in the central part of the Sierra del Chacaico (Cactus, Cheese, Boss and Sundial Gorges; Fig. 3 and the Sequence stratigraphy section later) have a higher proportion of tidalflat facies as a function of being furthest from points of fluvial input. This maximum distance from the centre of the tidal-channel meander belt translates directly into the zone of maximum preservation potential for tidal-flat deposits. Modern analogues such as the German Bight are prograding at rates that are much higher because of the sediment baffling effects of sea grasses (e.g. Biggs & Howell 1984; Zhuang & Chappell 1991), which did not evolve until the Eocene.

Sequence stratigraphy The basic sequence stratigraphic model for marine basins assumes that accommodation changes are driven by a combination of tectonic

Fig. 12. Holistic facies model showing the distribution of Lajas Formation facies associations and their spatial interrelationship. It is important to note that this is a freeze-frame and that delta-front deposits will be reworked during progradation and tidal flats may overlie extensive tidal-channel meander belts.

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subsidence and eustatic sea-level change (Vail et al 1977; Posamentier & Vail 1988). Relative falls in sea level may expose all or part of the shelf and, as the fluvial profile adjusts to the lowered base level, the rivers of the previous highstand are generally deepened and widened, resulting in incised valleys. During subsequent relative sea-level rise, these valleys are flooded and converted to estuaries. Commonly, the topographic restriction of the estuary 'funnel' provides sufficient amplification of the local tidal wave to allow deposition under meso- to macrotidal conditions (Dairymple et al. 1992; Zaitlin et al. 1994). The estuarine succession is transgressive and, once the valley is filled, deposition usually reverts to non-tidal open coastline shoreface or mouthbar facies associations of late transgressive systems tracts (TST) and highstand systems tracts (HST) (Zaitlin et al. 1994). The key point, with respect to tidal processes and products in this model, is that they are temporally restricted to times of early base-level rise (late LST/early TST) and spatially restricted to estuaries within incised valleys. In high subsidence settings (typically extensional basins and growth-faulted areas), the rate of hanging-wall subsidence can equal or exceed the rate of eustatic fall, resulting in no relative fall in sea level (Gawthorpe et al. 1994; Ho well & Flint 1996). In such cases, the sequence

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boundary may be identified by a change in facies stacking patterns and need not necessarily be associated with subaerial exposure and valley incision of former shelf deposits (Posamentier & Vail 1988). Sequences are asymmetric and dominated by the rise component of the sea-level cycle, producing thick, aggradational shallowmarine successions (e.g. Howell et al. 1996; Ravnas & Steel 1998). Lajas Formation facies associations exhibit complex spatial and temporal relationships, but are arranged in systematic packages bounded by regionally correlatable surfaces. A variety of these packages and their bounding surfaces are considered within a sequence stratigraphic framework below. Flooding surfaces Within the Lajas Formation, certain marked changes in grain size and intensively bioturbated Glossifungites surfaces can be traced over the length of the outcrop belt. Across these surfaces there are abrupt changes in facies from shallower water (e.g. muddy intertidal-flat deposits) below, to deeper water (e.g. subtidal-flat sandstones) above (e.g. Fig. 14). The successions between these surfaces are discussed in the next section. Surfaces that separate older strata below from younger strata above and across which there is

Fig. 13. Oblique aerial photograph of a typical Lajas cliff line with the parasequence and sequence boundaries highlighted. Note the low-order cycles from sand- to mud-rich deposition. The vertical scale is non-linear owing to parallax effects, but the whole succession is c. 500 m thick. Vertical lines represent faults, which are typically subparallel to the outcrop.

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Fig. 14. Typical stratigraphic section taken from the study area in Figure 5, showing example expressions of surfaces of stratigraphic importance that can be traced throughout the study area: (A) a sequence boundary overlain by tidal channel sandstones with thick mud-draped bedding surfaces. The mudstone horizons are interpreted as fluid mud deposits formed in the mixing zone of an estuary (incised valley fill); (B) tidal-channel sandstone with abundant Y-branching trace fossils of Polykladichnus isp. that originate from a marine flooding surface in a soft-ground expression of a marine flooding surface (i.e. not a classic Glossifungites surface, cf. MacEachern et al. 1992); (C) marine flooding surface represented by a marked change in lithology from tidal-flat facies to offshore shale facies.

evidence of an abrupt increase in water depth are defined as flooding surfaces (Van Wagoner et al. 1988). In non-tide-dominated systems, recognition of flooding surfaces is straightforward. Demonstrably marine shales typically overlie proximal mouthbar, upper shoreface, beach or delta-top deposits. Such expressions of flooding surfaces also occur in the coarsening-upwards tidal-delta deposits of the Cura Niyeu Formation (Fig. 13). However, most Lajas Formation facies were deposited within the intertidal range, with a

maximum interpreted bathymetry of 7-8 m. Coupled with the dominantly aggradational stacking pattern (see below) and a high degree of lateral variability within facies belts owing to tidal-channel erosion and migration, forced facies shifts due to relative sea-level rise are subtle. In a single vertical section (outcrop or well log) the recognition of true flooding surfaces is further complicated because the trace fossil signature of flooding surfaces can be mimicked by that found in laterally restricted, bioturbated,

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muddy tidal-channel abandonment deposits. Despite these problems, the extensive outcrop control allows regional flooding surfaces to be distinguished from local occurrences of similar areally restricted facies (e.g. FA 2c). In many cases, Lajas Formation flooding surfaces occur as an abrupt grain-size increase rather than the more typical abrupt grain-size decrease expected across flooding surfaces in shoreface successions. This relationship is typified by subtidal (deeper water) sandstones overlying, intertidal (shallower water) heterolithic strata. Parasequences Parasequences are defined as conformable sets of beds, separated by marine flooding surfaces or their correlative surfaces (Van Wagoner et al. 1988). In storm/wave-dominated settings, such as the classic Cretaceous Western Interior Seaway, USA, parasequences are upwardscoarsening (progradational) units with sheetlike, coast-parallel geometries; they shale out down-dip and pass up-dip into coastal-plain strata (Van Wagoner et al. 1990; Kamola & Van Wagoner 1995; Howell & Flint 2002). Fluvial-dominated deltas produce broadly similar parasequences, which are commonly lobate with more complex shale intercalations (e.g. Pulham 1989; Bhattacharya & Walker 1992). Using the flooding surface framework described above, intervening genetically related sets of beds bounded by flooding surfaces in the Lajas can be interpreted as parasequences. Unlike the uniform character of parasequences in storm/ wave- and fluvial-dominated successions, the distribution of facies within these systems is considerably more complex (Mcllroy et al. 1999). This complexity arises partially from the high degree of lateral facies variation that inherently exists within tide-dominated depositional systems and partially because of the potential for significant erosion within tidal systems. The typical thickness of a parasequence (approximately 10-15 m), a measure of the incremental relative sea-level rise that occurs between flooding surfaces, is comparable to documented depths of tidal scour (e.g. Yang & Nio 1989; Willis & Gabel 2001). Consequently, in contrast to wave- and river-dominated depositional systems in which parasequences contain a record of virtually continuous sedimentation with minor storm- or underflow-related erosion, tide-dominated parasequences are much more complicated and variable, containing multiple autocyclically controlled - phases of erosion and deposition. Three main types of parasequence succession occur between the flooding surfaces

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in the Lajas Formation: tidal deltas; tidalchannel sandstone complexes; and heterolithic parasequences characterized by prograding tidal-flat successions that may pass laterally into bayfill successions. Tidal deltaic successions (FAs la and Ib). These flooding-surf ace bounded, upwardscoarsening progradational units are up to 20 m thick and resemble classic parasequences of wave- and river-dominated deltas (Van Wagoner et al. 1990), except for the dominance of tide-generated internal structures (Fig. 15a). Tidal delta parasequences can be traced for over 18 km (Fig. 13) and are characterized by seaward - or obliquely - dipping clinoform geometries with the tops of the upwards-coarsening units commonly cut by genetically related single-storey channels (Fig. 6b). The greater thickness of these parasequences is probably due to their progradation into accommodation in front of the previous shoreline, such that the 20 m is comprised partially of the sea-level rise and partially of the older space, incompletely filled by the last progradation. Tidal-channel-dominated successions (FAs 2a-2c). These successions are sandstone dominated, comprising a variety of tidal-channel facies with subsidiary tidal-flat and prodelta/ bayfill facies. Muddy prodelta, bayfill or abandoned channel facies mark the basal flooding surface and are overlain by laterally extensive, multistorey tidal channel complexes (Fig. 15b). The upper portion of parasequences commonly contains more heterolithic channel-fill and tidal-flat facies resulting in a blocky to finingupwards grain-size profile. Away from the main tidal-channel belts, parasequences are likewise sand-rich, but are dominated by thick aggradational successions of tidal-flat facies. Heterolithic-dominated successions (FAs Ic and 3a-3c). These flooding-surf ace bounded successions are dominated by heterolithic tidal-flat and bayfill mudstone facies with minor volumes of meandering tidal-channel deposits (Fig. 8a). The vertical profile is highly variable and may show blocky, coarsening-upwards or fining-upwards motifs that reflect the coastparallel passage from tidal flats into large sediments-starved coastal embayments. Where tidal-flat dominated, the upper parts of parasequences commonly contain single-storey meandering tidal-channel facies (Fig. 15c), which form planar sand bodies up to 500 m wide. Some vertical profiles resemble the idealized

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Fig. 15. Example parasequences from the Lajas Formation showing the development of the parasequences bound by marine flooding surfaces, the proportion of mud to sand and the orientation of heterogeneities. The block diagrams attempt to represent some of the mesoscale facies variability and the logs are idealized to give an impression of a typical expression in core. In all cases the sections are hypothetical in nature, although based on a sound understanding of mesoscale facies variability as previously published (Mcllroy et al. 1999; Brandsaeter et al. 2005): (A) sand-rich tida channel-dominated parasequence in which the earliest deposited sediments (tidal-flat or bayfill deposits) are cut by multistorey tidal-channel deposits and capped by intertidal flat deposits (these are commonly eroded if overlain by another tidal-channel-dominated parasequence; (B) heterolith-dominated tidal-flat-rich parasequence showing a progradational succession from subtidal-flat facies through intertidal-sandflat facies with tidal creeks to a more heterolithic upper portion that may have tidal-channel sands concentrated at the top; (C) idealized tidal-delta parasequence showing a simple upwards-coarsening profile from offshore mudstones through mouthbars deposits that are typically cut by tidal-channel sandstones. Associated facies are delta-front turbidites and large-scale slump structures in the delta-front deposits.

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fining-upwards tidal-flat parasequence of Van Wagoner et al (1990).

the poorly exposed Mano sequence not discussed herein.

Sequence boundaries In addition to parasequence bounding flooding surfaces, the succession is also punctuated by four, regionally extensive, surfaces across which shallower water facies are abruptly juxtaposed on deeper water deposits. These surfaces are typically irregular, with topography resulting from incision into the underlying units. A maximum depth of erosion of 30 m was observed. Typically, the erosion takes the form of broad, wide scours into offshore marine shale, tidal mouthbar, tidal-flat and bayfill deposits. Overlying the surfaces are thin lenses of fluvial gravel, stacked, multistorey tidalchannel complexes and, more rarely on topographically high interscour areas, palaeosols are observed (e.g. Fig lie). In many cases parasequences onlap the incision surfaces (Fig. 5). These surfaces are interpreted as sequence boundaries. Recognition criteria for type 1 sequence boundaries in shallow-marine successions have been summarized by Van Wagoner et al (1988, 1990), and include a regionally developed subaerial erosion and/or exposure surface. The facies juxtaposition across this surface must reflect a basinward shift in facies, typically marked by tidal-channel, tidal-flat, fluvial or mature palaeosol facies lying sharply on lower shoreface or prodelta strata. The interpretation of sequence boundaries in the Lajas is based on the lateral extent of the surfaces (greater than any single facies belt), maximum depth of incision (deeper than documented tidal-channel scour), the subsequent onlap by more than one parasequence (e.g. Fig. 5), and, in certain cases, the association of fluvial deposits and or interfluve palaeosols. The sequences that these surfaces bound are described below. Unlike classic models, the sequence boundaries are associated with comparatively shallow scours (although still deeper than normal autocyclic tidal incision) rather than 'valleys' and, in most cases, facies changes remain within the tide-dominated facies belts.

Sequence 1 (Burgess Sequence) The lowermost sequence comprises a basal regional erosion surface that juxtaposes a 7-10 m-thick tidal channel complex on offshore marine shelf shales (Fig. 16). The tidal channel complex has a valley like geometry. The fill shows a deepening-upwards facies trend, indicative of base-level rise, and is interpreted as early TST. The valley fill is overlain by marine shelf shales and a complex succession of seven stacked parasequences comprising tide-dominated delta front and prodelta-offshore facies. The lower 15 m of this succession consists of a retrogradational parasequence set, with increasing proportions of FA la upwards, typical of a marine TST. The overlying 10 m of succession is aggradational and is overlain by a thick succession of rapidly prograding sand-dominated delta-front facies (FA Ib) and genetically linked tidal channels of FA 2a. The parasequence set is overlain by offshore mudstones (FA la) that are correlatable over the entire study area (Figs 13 & 16) and are interpreted as a maximum flooding surface. These seven parasequences are interpreted as a late transgressive systems tract and the unusual, strongly progradational parasequence contained within this TST is tentatively ascribed to a higher frequency minor relative sea-level fall within the background rise. Overlying the maximum flooding surface (base of the Paloma Parasequence, Fig. 16) is a strongly progradational set of tidally influenced delta parasequences, which is interpreted as a highstand systems tract. The total thickness of the Burgess Sequence is 110 m.

Evolution of the Cura Niyeu-Lajas sequence set The study interval has been subdivided into four sequences based on the recognition of lateral extensive sequence boundaries. These sequences are described in chronological order below, starting with the Burgess Sequence, which overlies

Sequence 2 (Pushme—Fullyou Sequence) The base of sequence 2 is an erosional sequence boundary that truncates the highstand tidally influenced deltas of sequence 1 with up to 30 m of incision (Fig. 16). Overlying the sequence boundary is a 35 m-thick aggradational stack of amalgamated tidal meander-belt deposits (early TST) that passes upwards into three heterolithic-dominated parasequences comprised predominantly of tidal-flat sediments. The Pushme -Pullyou Sequence is 140 m thick and shows a gross-scale aggradational to weakly retrogradational stacking pattern. The uppermost (Pratis) parasequence is particularly argillaceous. We interpret this whole succession as deposited during long-term slow baselevel rise when sediment supply was

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approximately balanced by the rate of accommodation generation.

Sequence 3 (Komplott Sequence) The basal Komplott sequence boundary exhibits up to 20 m of incision into the muddy heterolithic strata of the underlying parasequence, with a shallow, wide valley geometry. The internal facies architecture of the incised-valley fill is documented in Figure 5 and comprises two tidal-channel parasequences separated by a valley-wide marine flooding surface. Local palaeosols are developed on the interfluve surface. The Komplott Sequence is 110 m thick and shows a similar fining-upwards trend to sequence 2 (Fig. 16). The upper part of this succession includes the Norwegian Parasequence, which was described in detail by Mcllroy et al (1999).

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Two possible sequence stratigraphic interpretations are: (1) the entire succession represents a series of stacked incised-valley fills, similar to the interpretation forwarded by Van Wagoner (1991) for the tide-dominated Cretaceous Sego Sandstone in Utah; and (2) the tide-dominated, deltaic system maintained a tidal character throughout complete base-level cycles, with tidal sedimentation not confined to incised valleys. On the balance of field evidence, the latter interpretation is proposed for the Lajas Formation and the maintenance of tidal conditions is considered a function of the long-lived, structurally controlled embayment of the southern Neuquen Basin. This interpretation is similar for the aggradational tidal successions of the Tilje and He formations, offshore mid Norway (Martinius et al. 2000; Mcllroy 2004), the penecontemporaneous Neill Klinter Formation of Greenland (Dam & Surlyk 1995) and the Jurassic of the Inner Hebrides, Scotland (Mellere & Steel 1996).

Sequence 4 (Owl Sequence) The uppermost Lajas sequence is 180m thick. The basal sequence boundary almost completely truncates the uppermost (Norwegian) parasequence of sequence 3, with up to 25 m of incision (Fig. 16). The lower quarter of the sequence is dominated by tidal-channel fills and the upper section comprises heterolithic-dominated parasequences. The up-dip sections to the south (Fig. 16) show an upwards passage into fluvial channel sandstones, which supports a progradational trend in the upper two or three parasequences. However, the upper part of this sequence is poorly exposed, making detailed sequence stratigraphic analysis difficult. The sequence is overlain by fluvial deposits of the Challaco Formation (Fig. 16). The Lajas sequence stratigraphic model for meso-/macrotidal deposits Some shallow-marine systems are tidal in character over hundreds of metres in thickness.

Lowstands The occurrence of shallow incised-valleys implies periods of negative accommodation on the Lajas coastal plain; consequently, the sediment removed and bypassed from the shelf must have been deposited somewhere during lowstand times. Previous workers have interpreted broadly time-equivalent turbidites to the north as lowstand deposits associated with sequence boundaries in the Lajas (Gulisano & Gutierrez Pleimling 1994). However, it is not possible to physically trace the stratigraphy between the sections owing to incomplete exposure and structural complexity. Petrographic studies show that the turbidites are dominated by volcanic arc material (Eppinger & Rosenfeld 1996). Palaeocurrent data from outcrops also support their derivation from the early Andean arc to the west, rather than the Lajas shelf to the south (Burgess et al. 2000). The nature of the associated lowstand deposits therefore remains to be demonstrated.

Fig. 16. Summary diagram showing the Lajas Formation sequence set in the Sierra de Chacaico region. Facies associations are shown and both the sequence boundaries and parasequence boundaries highlighted along with the location of the logged section shown in Figure 5 and the cliff line in Figure 13. The concentration of tidal channel meander belt facies at the bases of parasequences and the gradual northward progradation of the fluvial Challaco Formation is noteworthy as is the limited lateral extent of the tide-dominated deltas of the basal Lajas (Cura Niyeu) Formation. Most of the parasequence boundaries have been walked out in the field. Trace fossil abbreviations: Arenic., Arenicolites isp.; Asteria., Asteriacites ispp; Astero., Asterosoma isp.; Chond., Chondrites isp.; Crab., ?crab burrows; Dacty., Dactyloidites isp.; Didy., Didymaulichnus lyelli; Diplo., Diplocmterion parallelum; Ophio., Ophiomorpha ispp.; Parah., Parahaentzschelinia isp.; Pphy, Palaeophycus isp.; Phoeb., Phoebichnus isp.; PlanoL, Planolites ispp.; Protov., Protovirgularia isp.; Rhiz., Rhizocorallium isp.; Ross., Rosselia isp.; Scoli., Scolicia isp.; Siphon., Siphonichnus isp.; Skol., Skolithos isp.; Tereb., Schaubcylindrichnus isp.; Teich., Teichichnus isp.; Thai, Thalassinoides isp.

D. McILROYCTAL.

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Base-lev el rise The stacked sandy tidal channel complexes above Lajas sequence boundaries are interpreted to have been deposited during a slow rate of accommodation creation resulting from baselevel rise (Fig. 16). With slow, continuous accommodation generation, meandering tidal channels extensively rework the tide-dominated coastal plain, producing wide meander belts and little or no preserved tidal-flat deposits (e.g. the Bert, N.W.C. and Noqui parasequences, see Fig. 16). Therefore, early transgressive systems tracts in tide-dominated systems are dominated by tidal-channels deposits. The sand-rich stacked tidal channel belts are overlain by heterolithic single-storey tidalchannel fills and more extensive tidal-flatdominated parasequences. We interpret this transition as reflecting increasing rates of relative sea-level rise. During periods of more rapid accommodation generation, meander belts will be narrower and show less of a tendency towards amalgamation (e.g. the Familiar, Perro Loco and Pratis parasequences, see Fig. 16). Lateral to tidal-channel meander belts, extensive tidal flats developed, with sediment being transported by longshore processes (e.g. the Skull Parasequence between Cactus Gorge and Sundial Gorge, see Fig. 16). Where longshore sediment supply is limited, especially during periods of rapid accommodation generation, mud-dominated coastal embayments develop (e.g. much of the Pratis Parasequence between Armadillo Gorge and Cheese Gorge, Fig. 16). Consequently, the later part of sequences in tide-dominated systems, such as the Lajas, will be characterized by rapidly prograding tidal-flat deposits, and tidal channels will be preserved as isolated bodies. The deposits of restricted coastal embayments may also be common. Maximum flooding surfaces are difficult to identify precisely due to the limited bathymetry range but 'zones' can be identified where the retrogradational muddy tidal-channel/tidal-flat facies above the stacked tidal channel complexes pass upwards into aggradational-weakly progradational sets of laterally extensive tidal-flat parasequences. These can be tentatively interpreted as highstand systems tract deposits and there are two important implications for this model: Tidal highstand systems tracts are much muddier than in conventional shoreface or deltaic systems and are therefore of low reservoir quality.

Conventional sequence stratigraphic analysis from well logs (and core, if not correctly interpreted) would interpret the Lajas Formation as comprising only transgressive systems tracts.

Conclusions Completefining-upwardstidal-flat successions are rare in published descriptions of ancient tidal successions because in estuarine settings tidal flats are continuously destroyed by tidal-channel migration across the width of the estuary. In contrast, completely preserved tidal-flat systems are common in the Lajas Formation, owing to the extensive progradational-aggradational tidal coastline setting, in which tidal-channel meander belts were unable to migrate across the whole coastline. In large, structurally defined seaways and embayments, tidal sediments are not restricted to incised-valley fills if the geometry of the basin is such that it can amplify tidal currents over long periods, irrespective of base level (Mcllroy et al 1999; Willis et al. 1999; Bhattacharya & Willis 2001; Willis & Gabel 2002). If this criterion is met in an actively subsiding basin, great thicknesses of tidal sediments may be deposited along tidal coastlines, with both regressive and transgressive trends, depending on the sediment supply to accommodation balance. Lajas flooding surfaces are marked by a range of 'deeper on shallower' facies juxtapositions that include a fine to coarse grainsize shift where subtidal sandstones overlie intertidal-flat mudstones. The flooding surfaces delineate three types of parasequence that exhibit coarsening-up wards, finingupwards and coarsening- to fining-upwards motifs. Sequence boundaries do not show deep incision and no definite lowstand deposits are yet identified. Stacked, multistorey tidal-channel complexes overlying sequences boundaries reflect low rates of accommodation creation during early base-level rise. An accelerating rate of base-level rise and space creation is marked by the upwards change to isolated heterolithic tidal-channel fills and extensive tidal-flat- and bayfilldominated parasequences. Large rates of accommodation generation may also account for the suppressed incision at sequence boundaries.

TIDAL SEQUENCE STRATIGRAPHY We wish to acknowledge the financial support of Statoil, Neste, Saga and Norsk Conoco. Statoil and Saga staff, especially A. Martinius, A. Naess, C. Jourdan, K. Gibbons, S. Str0mmen, P. Ringrose, G. Saigal and R. Knarud, are thanked for their support and thoughtprovoking discussions. Assistance with field logistics and access to subsurface data was provided by YPF s.a. and helpful discussions with N. Bolatti, I. Brisson, R. Veiga and R. Manoni. S. Flint would also like to recognize the help of M. Uliana, now deceased, and our introduction to the Neuquen Basin by L. Spalletti of La Plata University. This paper has been much improved by the constructive comments of J. Bhattacharya, L. Buatois, C. Fielding & B. Willis.

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River, Western Australia. Journal of Geology, 86, 621-642. DALRYMPLE, R.W., ZAITLIN, B.A. & BOYD, R. 1992. Estuarine facies models: conceptual basis and stratigraphic implications. Journal of Sedimentary Petrology, 62, 1130-1146 DAM, G. & SURLYK, F. 1995. Sequence stratigraphic correlation of Lower Jurassic shallow marine and paralic successions across the Greenland-Norway seaway. In: STEEL, R.J., FELT, V.L., JOHANNESSEN, E.P & MATHIEU, C. (eds) Sequence Stratigra phy on the Northwest European Margin. NPF, Special Publications, 5, 483-509. DEAN, J.S. 1987. Depositional environments and paleogeography of the lower to middle Jurassic Cuyan Group, Neuquen Basin, Argentina. DPhil References Thesis, Colorado School of Mines. ALLEN, G.P. 1991. Sedimentary Processes and facies DE BOER, P.L., COST, A.P. & VISSER, M.J. 1989. The diurnal inequality of the tide as a parameter for in the Gironde Estuary: a recent model for macrorecognizing tidal influences. Journal of Sedimentidal estuarine systems. In: SMITH, D.G., tary Petrology, 59, 912-921. RENISON, G.E., ZAITLIN, B.A. & RAHMANI, R.A. (eds) Clastic Tidal Sedimentology. Canadian DE MOWBRAY, T. & VISSER, M.J. 1984. Reactivation surfaces in subtidal channel deposits, OosterSociety of Petroleum Geologists Memoirs, 16, schelde, Southwest Netherlands. Journal of Sedi29-40. mentary Petrology, 54, 811-824. ALONGI, D.M. 1991. The role of intertidal mudbanks in the diagenesis and export of dissolved and par- DIGREGORIO, J.H. & ULIANA, M.A. 1980. Cuenca neuquina. Segundo Simposio de Geologia Regional ticulate materials from the Fly Delta, Papua New Argentina, Academia Nacional de Ciencias, Guinea. Journal of Experimental Marine Biology Cordoba, 2, 985-1032. and Ecology, 149, 81-107. BAKER, E.K., HARRIS, P.T., KEENE, J.B. & SHORT, EPPINGER, K.J. & ROSENFELD, U. 1996. Western Margin and provenance of sediments of the S.A. 1995. Patterns of sedimentation in the macroNeuquen Basin (Argentina) in the Late Jurassic tidal Fly River delta, Papua New Guinea. In: and Early Cretaceous. Tectonophysics, 259, FLEMMING, B.W. & BARTHOLOMA, A. (eds) Tidal 229-244. Signatures in Modern and Ancient Sediments. International Association of Sedimentologists, FEMES, H., DE RESSEGUIER, A. & TASTET, J.-P. 1999 Intertidal clay-drape couplets (Gironde estuary, Special Publications, 24, 193-211. France). Sedimentology, 46, 1-15. BHATTACHARYA, J.P. & WALKER, R.G. 1992. Deltas. In: WALKER, R.G. & JAMES, N.P. (eds) Facies FREY, R.W., HOWARD, J.D., HONG, J.-S. & PARK, B.-K. 1989. Sediments and sedimentary sequences Models - Response to Sea-level Change. Geologion a modern macrotidal flat, Inchon, Korea. cal Association of Canada, 157-179. Journal of Sedimentary Petrology, 59, 28—44. BHATTACHARYA, J.P. & WILLIS, B.P. 2001 Lowstand deltas in the Frontier Formation. Powder River GAWTHORPE, R.L., FRASER, A.J & COLLIER, R.E. 1994. Sequence stratigraphy in active extensional basin, Wyoming: implications for sequence stratibasins: implications for the interpretation of graphic models. AAPG Bulletin, 85, 261-294. ancient basin-fills. Marine and Petroleum BIGGS, R.B. & HOWELL, B.A. 1984. The estuary as a Geology, 11, 642-658. sediment trap: alternate approaches to estimating its filtering efficiency. In: KENNEDY, V.S. (ed.) GULISANO, C.A. & GUTIERREZ PLEIMLING, A.R. 1994. Neuquen Basin, Neuquen Province. Field The Estuary as a Filter. Academic Press, Trip Guidebook A. In: Proceedings of the 4th InterNew York, 107-129. national Congress on Jurassic Stratigraphy and BRANDS^TER, L, MC!LROY, D., LIA, O. & Geology. Secretaria de Mineria de la Nacion, RINGROSE, P. 2005. Integrated modelling of Lajas Buenos Aires, Argentina, 1-111. Formation tide-dominated deltas. Petroleum HARRIS, P.T., BAKER, E.K., COLE, A.R. & Geoscience, 11, 37-46. SHORT, S.A. 1993. Preliminary study of sedimenBRIDGES, P.H. & LEEDER, M.R. 1976. Sedimentary tation in the tidally dominated Fly River Delta, model for intertidal mudflat channels with Gulf of Papua. Continental Shelf Research, 13, examples from the Solway Firth, Scotland. Sedi441-472. mentology, 23, 533-552. BURGESS, P.M., FLINT, S. & JOHNSON, S. 2000. HOWELL, J., FLINT, S. & HUNT, C. 1996. Sequence stratigraphic signatures in a complex shelf succesSequence stratigraphic interpretation of turbiditic sion: the Fulmar sands of the North Sea Central strata: an example from Jurassic strata of the Graben. Sedimentology, 43, 89-114. Neuquen basin, Argentina. Geological Society of HOWELL, J.A. & FLINT, S. 1996. A model for high resAmerica, Bulletin, 112, 1650-1666. olution sequence stratigraphy within extensional COLEMAN, J.M. & WRIGHT, L.D. 1978. Sedimentation basins. In: AITKEN, J.F. & HOWELL, J.A. (eds) in an arid macrotidal alluvial river system: Ord

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Sedimentary evolution and depositional architecture of a lowstand sequence set: the Lower Cretaceous Mulichinco Formation, Neuquen Basin, Argentina ERNESTO SCHWARZ1'2 & JOHN A. HOWELL3'4 Centro de Investigaciones Geologicas, University of La Plata-CONICET, La Plata B1900TAC, Argentina 2 Present address: Department of Earth Sciences, University of Ottawa, 140 Louis Pasteur St, Canada KIN 6N5 Ottawa, Ontario (e-mail: eschwarz@ uottawa. ca) 3 STRAT Group, Department of Earth Sciences, University of Liverpool, Liverpool L69 3GP, UK 4 Present address: Centre of Integrated Petroleum Research, University of Bergen, Allegaten 41, N-5007 Bergen, Norway l

Abstract: The Valanginian-aged Mulichinco Formation was deposited in the Neuquen Basin (west-central Argentina) during and immediately after a major fall in sea level, partially triggered by a tectonic inversion pulse. The formation represents a lowstand wedge where excellent outcrops, together with refined biostratigraphic coverage, have permitted the detailed examination of contemporaneous non-marine and marine deposits. Fourteen facies associations were identified in the Mulichinco Formation. They represent accumulation in a variety of environments ranging from gravelly fluvial braidplains to outer-shelf marine settings. Distribution of depositional environments, together with the identification of key surfaces and stratal patterns, has resulted in the identification of early and late lowstand, transgressive and highstand systems tracts. Accordingly, the Mulichinco lowstand wedge comprises one third-order sequence that lasted about 2 Ma and represents a lowstand sequence set. The character of shoreline sedimentation was highly variable along strike within the Mulichinco depositional area and alluvial deposits were not developed within incised valleys. Tectonically derived topography, basin physiography and fault-controlled subsidence are interpreted to have been the main controls on the evolution of the Mulichinco lowstand wedge. The results of this study have important implications for understanding both the history of Neuquen Basin and illustrating the previously undocumented architectural complexity that may exist within lowstand wedges.

Within the classic sequence stratigraphic models the lowstand systems tract occurs as either basinfloor fans, if the relative sea-level fall exposes a shelf-slope break, or as lowstand shorelines, if the basin has a ramp-style physiography or the shelf-slope break is not subaerially exposed (Posamentier & Vail 1988; Van Wagoner et al. 1988). Existing models for ramp settings imply that incised valleys cut during falling sea level fed sediment to forced-regressive shoreface and deltaic systems. Incised valleys are then filled with alluvial sediments during the late lowstand and early transgressive systems tracts (TST) (Hettinger et al. 1994). However, lowstand systems tracts can show greater variability than is predicted from the standard models. Ramp-type

basin physiography coupled with tectonics can interplay to produce a more complicated lowstand response (e.g. Posamentier & Allen 1993). Moreover, the increased amplification of tidal currents by topographic irregularities during lowstand intervals (Nummedal et al 1993) is also potentially important in tectonically active basins (Howell & Flint 1996). This study documents in detail the facies associations and stratal architecture of the nonmarine-marine Mulichinco lowstand wedge, Superb exposures and a well-refined biostratigraphic scheme allow reliable correlation of 15 measured localities. Interpretation of facies associations, together with the identification of key surfaces and stratal patterns, has resulted in

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 109-138. 0305-8719/057$ 15.00 © The Geological Society of London 2005.

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the recognition of linked systems tracts within the lowstand wedge. Spatial and temporal distribution of the subenvironments in the Mulichinco lowstand wedge show a high degree of variability, and some common lowstand features such as incised valleys were not observed. The Mulichinco Formation was developed on top of a previous ramp setting (Legarreta & Uliana 1991), following a major fall in relative sea level. Mulichinco accumulation occurred immediately after a tectonic inversion pulse and accommodation was created by fault-controlled subsidence (Vergani et al 1995). Therefore, the Mulichinco Formation is an excellent example of a complex lowstand wedge developed within a tectonically active basin with a ramp-type physiography, and provides a unique opportunity to explore variability within lowstand wedge systems. The main objectives of this study are: (1) to document the sedimentary facies associations and stratigraphic relationships within the Mulichinco Formation; (2) to present a reliable sequence stratigraphic framework and palaeogeographic reconstructions for the unit; and (3) to consider the controlling factors on the evolution of the Mulichinco lowstand wedge and its implications for the better understanding of variability within lowstand systems tracts.

Geological setting and stratigraphy From the Middle Jurassic to Early Cretaceous the Neuquen Basin (west-central part of Argentina, Fig. 1) was a back-arc basin that lay inland of an emergent magmatic arc, which was associated with the eastward subduction of proto-Pacific oceanic crust beneath the western margin of Gondwana. The tectonic history during that time reflects regional thermal subsidence, interrupted by several episodes of structural inversion (Vergani et al 1995). In the Late Cretaceous the basin evolved into a foreland system that persisted until the final infilling during Early Tertiary times. The present-day area includes two different regions: an Andean fold and thrust belt to the west; and a hydrocarbonbearing platform to the east (Fig. 1). The Valanginian Mulichinco Formation belongs to the Mendoza Group (TithonianBarremian in age). From the Tithonian to the Early Valanginian an extensive second-order highstand developed, which includes offshore to continental deposits (Gulisano et al. 1984; Legarreta & Gulisano 1989) (Fig. 2). During that interval the depocentre was a semi-enclosed basin with a funnel-shaped morphology, open to the north and west through gaps in the arc (Fig. 3). The eastern and southern margins of

Fig. 1. Location map of the Neuquen Basin and the study area of the Mulichinco Formation.

the basin were characterized by a ramp profile (Legarreta & Uliana 1991; Spalletti et al 2000). A pulse of tectonic inversion in middle Early Valanginian times (Vergani et al 1995) partially (if not totally) accounts for the relative drop in sea level that led to the deposition of the Mulichinco Formation (Fig. 2). Vergani et al (1995) suggested that fault-controlled subsidence also occurred during the Early Valanginian. Consequently, the Mulichinco Formation was deposited at a time of differential tectonic subsidence, inversion and changing basin configuration. These factors potentially have a significant control on the facies distribution and depositional environments within a sequence stratigraphic framework. The 200-400 m-thick deposits of the Mulichinco Formation are largely restricted to the centre of the basin (Fig. 2). The base of the unit is a regional unconformity (the IntraValanginian unconformity of Gulisano et al 1984) that overlies black shales in the majority of the area (Vaca Muerta Formation). The top

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Fig. 2. Stratigraphic chart (after Legarreta & Gulisano 1989), second-order sequence stratigraphy and tectonic history (after Vergani et al. 1995) of the Berriasian-Early Hauterivian in the Neuquen Basin (modified from Schwarz 2003). Time scale after Gradstein et al. (1995). Note the wedge disposition of the study Mulichinco Formation.

of the Mulichinco Formation corresponds to a long-term transgressive phase that led to the basin-wide deposition of the euxinic black shales of the Agrio Formation (Fig. 2). A highly refined biostratigraphic scheme based on ammonites was established for the Early Cretaceous of the Neuquen Basin (Aguirre-Urreta & Rawson 1997, 1999) (Fig. 4). It dates the deposits of the Mulichinco Formation as middle Lower Valanginian-middle Upper Valanginian (c. 2 Ma, Fig. 4), and provides a very well constrained biostratigraphic control upon correlations within the marine portion of the Mulichinco Formation. Study area and methodology The study area is located in the deformed Andean region of the basin, between 37° and 38°20'S

latitude (Fig. 1). The excellent present-day exposures of Mulichinco Formation crop out along the flanks of a series of N-S-trending anticlines (Fig. 5). The majority of the outcrops are steeply dipping and laterally extensive for several kilometres. The Mulichinco Formation was studied in 15 localities that cover the entire exposures of the unit (Fig. 5). In nine of the localities the complete Stratigraphic interval is exposed, whereas in the other six at least half of it crops out. Spacing between the study localities is between 5 and 30 km. When considering the distribution of the sediments it is useful to group these localities into three sectors (the Southern, Central and Northern regions in Fig. 5). Sections were logged at 1:100 using standard logging techniques. Texture, sedimentary structures, sediment body thickness, geometry and

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and X-ray diffraction studies of selected samples were conducted at the laboratory.

Fades associations

Fig. 3. Schematic palaeogeography of the TithonianBerriasian Neuquen Basin, and its relationship with the Andean magmatic arc (modified from Spalletti et al. 2000). The basin was characterized by a ramp profile in the southern and eastern margins, and was located approximately 10° further south than it is at the present day.

Within this study, the Mulichinco Formation has been subdivided into 14 facies associations (FAs), each of which represent recurring combinations of facies that are, broadly, environmentally related (Reading & Levell 1996). Details of the FAs are presented in Table 1 and their distribution along the study area is shown in representative logs in Figure 6. Of the facies associations, 13 are totally or predominantly siliciclastic while only one is a fully carbonate association (FA 14). The siliciclastic FAs are interpreted to represent fluvial braidplains (FAs 1 and 2), fluvio-dominated (FA 3) and tide-influenced (FAs 4 and 5) deltaic systems, meandering fluvial plains and estuaries (FAs 6 and 7), lowenergy coastal plains and embayments (FAs 8 and 9), and open-marine shoreface-shelf environments (FAs 10-13). Facies associations representing fluvial braidplains are found only in the Southern Region, whilst the open marine FAs are the exclusive components of the Mulichinco Formation in the Northern Region (Fig. 6).

spatial relationships were documented during field work (Fig. 6). Palaeocurrents, as well as faunal and ichnofaunal content and taphonomy, were also recorded. Micropalaeontological, Fluvial braidplains facies associations content was analysed by S. Ballent at the FA 1: proximal braidplain. Rounded, granule Museum of La Plata (Argentina). Petrographic conglomerates interfingered with pebbly

Fig. 4. Valanginian biostratigraphic scheme of the Neuquen Basin (compiled from Aguirre-Urreta & Rawson 1997, 1999), and the temporal distribution of the Mulichinco Formation (after Schwarz 2003). Time scale after Gradstein etal (1995).

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Fig. 5. Simplified geological map of the study area, showing Mulichinco Formation outcrops, location of measured sections and the three regions in which they were grouped. BL, Barranca de Los Loros; CP, Casa de Piedra; CC, Cerro Curaco; ED, El Durazno; EC, Estancia Coihueco; LM, Los Menucos; LP, La Parva; LR, Loma Rayoso; PC, Puerta Curaco; PN, Pichi Neuquen; QT, Quintuco; RH, Rahueco; RS, Rio Salado; TR, Trahuncura; TE, Treleitube.

Table 1. Description and interpretation of fades associations recognized in the Mulichinco Formation FA

General organization

FA1

Lenticular units (0.7-3 m thick)

FA 2

Lenticular (Lt) (1-5 thick and tenshundreds of metres wide) and tabular (Tb) units

FA 3

General upwardscoarsening trend (50 m); tabular units at bottom (B) and middle (M); lenticular units at the top (T) (up to 3 m thick)

FA 4

Lenticular units (1.5 -7m thick and tenshundreds of metres wide)

Lithology and sedimentary structures

Bed/set thickness and internal organization

Granule conglomerates and pebbly sandstones; massive and planar cross-bedding prevail, minor imbrication and trough crossbedding Lt: pebbly-medium-grained sandstones; trough crossbedding dominant, but planar cross-bedding present; Tb: sandstones to siltstones; massive or with planar lamination; occasional mudcraks; smallscale channels intercalate B: laminated siltstones and micaceous sandstones with normal grading, flute marks, and rare asymmetric rippled tops; sandstone proportion 20-40% to top; M: medium- to coarsegrained sandstones; massive, planar or tangential crossstratified beds separated by thin shales and stacked vertically; convolute lamination; T: sandstones and pebbly sandstones; massive or with trough cross-bedding Fine- to coarse-grained sandstones; tangential cross-bedding dominant, but also sigmoidal and trough cross-bedding present; rippled tops and mud drapes common; convolute lamination

Decimetre- to 2 m-thick sets; weakly erosive bases; upwards-fining or trendless

Fossils and trace fossils

Occurrence

Interpretation

Barren

Southern Region; basal Mulichinco Formation

Low-sinuosity gravelly channels in proximal braidplain

Lt: sets <3 m; erosive and concave-up bases; upward- fining or trendless; Tb: beds <0.4 m thick; sharp bases and tops; variable trends

Logs and rare plant debris

Southern and Central regions; basalmiddle Mulichinco Formation

Low-sinuosity sandy channels and floodplain in distal braidplain

B: 0.02-0.25 m-thick beds; M: dm-scale beds with sharp bases; individual upwards-coarsening and -thickening packages (27 m thick); T: dm-scale sets with erosive bases

Abundant plant debris; rare bioclasts

Central Region; basal Mulichinco Formation

Fluvio-dominated prodelta- delta front and distributaries in delta plain

Sets 0.1-0.7m thick; erosive bases; upwards fining and bed thinning

Tops burrowed or capped by bioclastic carbonates

Central Region; basal-middle Mulichinco Formation

Distributary tidal channels

FA 5

Individual upwardscoarsening packages (decimetres to 12 m thick); tabular units

FA 6

Interbedded lenticular (Lf) units (1-3 m thick) and tabular (Tb) units

FA 7

Mainly tabular units (<7 m thick)

FA 8

Mainly lenticular units (0.5-2.5 m thick and 50250 m wide)

B\ massive mudstones and laminated siltstones; M: veryfine- to fine-grained sandstones and mudstone drapes; wavy/ lenticular discontinuous beds with symmetrical/asymmetrical ripples; synaeresis cracks; sandstone proportion 25-70% to top; T: large-scale inclined strata with cross-stratified sandstones and mud drapes Lt: pebbly -fine-grained sandstones; trough crossbedding pass to ripple lamination; large-scale inclined strata common; mud clasts and drapes parallel to surfaces; Tb: mudstone/sandstone couplets with lenticular to wavy bedding and sandstone beds with planar or massive bedding separated by mud laminae; rare mudcraks Massive mudstones, heterolithic intervals with wave and currentripple lamination, and tangentially cross-stratified sandstones and skeletal calcareous sandstones (some with landward migration) Large-scale inclined heterolithic strata of mudstones -sandstones couplets with mud draped ripples or low-angle stratification; trough crossbedding rare; ferruginous gravels at bases

M: cm- scale thick beds with sharp bases and sharp or gradational tops; upwards-coarsening and -thickening trends; T: plano-convex forms; gradational bases and sharp, erosive tops

Abundant plant debris; Rare Cruziana ichnofacies (mainly Phycodes, Teichichnus, Bergaueria}; shell fragments

Central Region; not widespread

Tide-influenced prodelta- delta front

Lt: sets <0.40 m; erosive bases and locally concave-up with clear cut banks; upwards-fining; Tb: cm to few dm thick beds; sharp bases and tops, commonly upwards-fining trends

Lt: logs and plant debris; Tb: rare gastropods and trace fossils (Palaeophycus and Teichichnus}

Southern and Central regions; middle Mulichinco deposits

Fluvial and tide-influenced meandering channels in fluvial plain and inner estuary with tidal flats

Sets 0.1-0.30 m thick, units with sharp bases and tops; either upwardsfining and -coarsening trends

Scarce marine ostracods+; common Cruziana ichnofacies

Same as FA 6

Central and outer estuary, and transgressive coasts

Beds cm- to a few decimetres thick; concave-up and erosive bases; upwards-fining trends

Planolites and Teichichnus burrows common, shell fragments

Southern Region; top Mulichinco Formation

Tidally influenced, suspended-load channels in a coastal plain

(continued}

Table 1. Continued FA

General organization

FA 9

Tabular units (< 10 m thick)

Massive dark shales and siltstones; less fine-grained sandstones with horizontal to wavy lamination; some sideritic nodules

Sandstones beds <0.1 m thick; upwardscoarsening trends

FA 10

Mainly tabular units (< 14 m thick)

Cross-stratified sets 0.2-0.7 m thick; sharp or erosive bases; upwards-coarsening or trendless

FA 11

Tabular units (<4 m thick)

Bioturbated fine- to niediumgrained sandstones with rare flat troughs and shell lags; also sandstones and oolithic/skeletal calcareous sandstones with trough and tangential crossbedding and abundant mud drapes passing laterally into wavy bedding; minor planar lamination and current ripples Very-fine- to fine-grained sandstones and oolithic calcareous sandstones; HCSdominated units with mainly symmetrical hummocks and ripple-dominated units with small-scale cross-lamination dipping in opposite directions

Lithology and sedimentary structures

Bed/set thickness and internal organization

Sets cm to a few decimetres thick; transitional bases; commonly upwardscoarsening and -thickening trends

Fossils and trace fossils Marine ostracods*; non-turritellid gastropods; rare Palaeophycus and Teichichnus; vascular plant fragments Common to rare Skolithos ichnofacies (mainy Ophiomorpha and Arenicolites}\ trigonid fragments; rare Thalassinoides Skolithos and Cruziana burrows common

Occurrence

Interpretation

Same as FA 8

Low-energy embayment, lagoon or shallow ponded water body

Northern -Central regions; top Mulichinco Formation

Upper -middle (subtidal) shoreface

Basal and top Mulichinco Formation in Northern Region; only at the top in Central Region

Storm-dominated to mixed tide- and storm-influenced lower shoreface

FA 12

Tabular units (<5.5m thick)

FA 13

Tabular units (< to 50 m thick)

FA 14

Tabular and very extensive units (20-50 m thick and up to several kilometres)

Interbedded siltstones and veryfine-grained sandstones; sandstone proportion 30-70% to top; micro-hummocks and combined flow-ripples dominant, but current ripples also present Massive mudstones, less marls; allochthonous shellbeds with thin, fine-grained (< 1 cm long) disarticulated fragments; parautochthonous shellbeds with thick beds (< 1.2 m) and gravelsize, mainly articulated bioclasts; also glauconite grains Marls, wackestones and densely packed oyster accumulations (boundstones), less skeletal packstones/rudstones; masssive beds dominant; some bioturbation; calcitic nodules common

0.01 -0.20 m-thick beds; sharp or transitional bases and gradational tops, upwards-coarsening trends

Frequent marine bivalves and burrows of Cruziana and Skolithos ichnofacies

Same as FA 1 1

Storm-dominated inner shelf

Sharp bases and transitional tops; trendless or upwards-coarsening

Abundant deep and shallow infaunal bivalves, ammonites, oysters; foramostracod association common Infaunal bivalves, cemented oysters (Ceratostreori), reclined oysters (Aetostreon), echinoids; radiolarianforam association

Same as FA 1 1

Outer shelf

Middle Mulichinco Formation in Northern Region

Clastic starvation and carbonate system (epiric ramp?) with oyster accumulations

Sharp bases and transitional tops, individual upwardscoarsening cycles (315m thick) or trendless

Lt, lenticular (channelized) units; Tb, tabular (non-channelized) units; cm: centimetre; B, bottom section of upwards coarsening package; M: middle section of upwards-coarsening package; T, top section of upwards-coarsening package. +Chyterella sp. *Cytherella sp., Cytherella montosaensis, Rostrocytheridea sp.

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sandstones characterize this facies association (Fig. 6; logs TR and QT, Table 1). Massive and planar cross-bedded facies are common, with rare pebble imbrication and trough crossbedding. Internally these units fine upwards and/or downcurrent to pebbly sandstones with trough or tangential cross-bedding (Fig. 7a). Units occur in lenticular-tabular bodies with sharp, weakly erosive basal surfaces. Finegrained deposits between the coarse-grained, lenticular bodies are rare to absent. Palaeocurrents are unimodal, showing a mean NE sediment transport direction (Fig. 8). This conglomeratic facies association is interpreted as the gradual infilling of active fluvial-channel complexes with transverse and longitudinal gravelly bars (Ramos & Sopena 1983). Continuous aggradation or gradual infilling of active channel complexes resulted in the deposition of the cross-bedded pebbly sandstones (Rust 1978). This facies association is interpreted to represent a proximal braidplain (sensu Rust & Koster 1984), where the conglomerate units accumulated under high sedimentsupply conditions. FA 1 represents the most source-proximal parts of the alluvial system, located towards the south and SW of the study area.

configuration is more likely to form in lowsinuosity, but relatively deep, channels, whose margins may be maintained by vegetation, mud and/or valley confinement (Rust & Gibling 1990). Avulsion was the main relocation mechanism in the alluvial plain, and it caused rapid abandonment of the active tract, so channel-fill sequences show abrupt tops and lack shallow-water components (cf. Rust & Gibling 1990). Fine-grained tabular deposits between main channels represent a relatively well drained overbank environment that was typically subaerially exposed. Silts were deposited from suspension during flooding, while the sandy tabular beds accumulated by ephemeral unconfined flows. Associated small-scale lenticular units are interpreted as crevasse channels (Clemente & Perez-Arlucea 1993). This facies association passes up depositional dip to FA 1 and both share a regionally consistent palaeoflow pattern (towards the NE), indicating that they were part of the same depositional system. The high proportion of channelized (as opposed to overbank) units, and apparent absence of levee deposits, suggest that this facies association most probably represents deposition from braided channels in a sandy distal braidplain (cf. Rust 1978; Browne & Flint 1994).

FA 2: distal braidplain. This facies association comprises lenticular (channelized) and tabular (non-channelized) units (Fig. 7b; Table 1). Lenticular units (l-5m thick) exhibit sharp and erosive concave bases commonly overlain by pebbles and/or rip-up clast horizons and silicified logs. These units are composed of pebblymedium-grained sandstones and exhibit a rather homogeneous internal fill dominated by trough cross-stratification (Fig. 6), but planar crossbedding is also locally present. The tops of the lenticular units are sharp and flat, but they can be eroded by overlying bodies. Palaeocurrent distribution is fairly unimodal toward the east and NE (Fig. 8). Subordinate tabular units are associated with the lenticular bodies (Fig. 7b). These units consist of medium-grained sandstones-siltstones (Table 1). Beds are typically tabular and massive or horizontally bedded, and mudcracks were seldom observed. Small-scale lenticular sandbodies (<1 m thick) are a minor component of these tabular units. In the Southern Region, the proportion and thickness of tabular units increase upwards (Fig. 6; logs TR and QT). Lenticular units are interpreted to represent the main channels of a fluvial system. Units dominated by trough cross-bedding indicate that three-dimensional (3D) dune fields occupied the major proportion of the channels. This

Deltaic facies associations FA 3: fluvio-dominated delta. This facies association is restricted to the basal portion of the Mulichinco Formation in the Central Region where it reaches 50 m in thickness (Fig. 6; PN section, 15-65 m). The succession exhibits a general upwards-coarsening trend and an increase in bed thickness (Table 1). In more detail, this broad pattern breaks down into stacked, metre-scale upwards-coarsening successions (Fig. 6). The lower part of the succession comprises rhythmic intervals with thinly bedded laminated siltstones/mudstones intercalated with fine-grained micaceous sandstones (Fig. 7c). Sandstones beds are normally graded and/or plane-parallel laminated with flute marks as common basal features. Plant debris is very abundant but fossils are absent. Middle parts of the facies association are characterized by massive, planar-laminated or tangential cross-bedded strata (Table 1). These decimetre-thick beds exhibit sharp basal surfaces and are separated by thin mudstones layers or sand-mud couplets. They stack vertically and laterally to build bar-like sandbodies. The upper portion of the facies association is composed of lenticular bodies with basal, highrelief surfaces that may erosionally overlie the

Fig. 6. Seected sedimentological ogs of Mulichico Formation within the stud area. The figure aso shows the facies associations and depositional environments identified in this stud and stratigraphic position of others ilustrations. See Figure 5 for og location

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lower portions of the successions (Fig. 7d). Bioclastic lags occur locally near the basal surfaces. These lenticular units consist of sandstone and pebbly sandstone beds, largely with trough cross-stratification (palaeotransport direction to the east). The regular upwards-coarsening trend and lithofacies of this facies association are suggestive of deposition in a deltaic setting (cf. Elliott 1986). Abundant plant debris and mica suggest

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proximity to a fluvial system (Martinsen 1990), while soft-sediment deformation and the absence of bioturbation may indicate high sediment rates (Bhattacharya & Walker 1991). Rhythmic intervals reflect alternating deposition from dilute effluent-derived flows and settling from suspension in a prodelta-distal mouth bar setting. More dense flows (turbidity currents?) generated graded sandy beds, most probably during exceptional flooding events (Martinsen

Fig. 7. Main characteristics of fluvial and deltaic facies associations, (a) Planar cross-bedded conglomerates (Gp) and pebbly sandstones (SGp) developed in proximal fluvial braidplains (FA 1). (b) Lenticular (Lt) and tabular (Tb) units of distal braidplain facies association (FA 2). (c) Thinly bedded, laminated siltstones and graded fine-grained sandstones that characterize the basal sector of the fluvio-dominated delta facies association (FA 3). (Lens cap diameter is 5 cm.) (d) Middle massive sandstones (Ms) and top lenticular units (Lt) of the same facies association (FA 3). (e) Tangential cross-bedded sandstones with abundant mud drapes and mud chips that fill distributary tidal channels (FA 4). (Coin diameter is 2.3 cm.) (f) Large-scale inclined strata, with transitional base and where sand-rich (S) and heterolithic-rich (H) intervals alternate, are interpreted as subtidal macroforms in the upper parts of tide-influenced delta fronts (FA 5). (Person for scale.)

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Fig. 8. Measured palaeoflow directions for the Mulichinco Formation. Palaeocurrent data has been grouped following facies association. Only data for selected localities are shown.

1990; Mellere et al 2002). Vertically and laterally stacked sandbodies composed of massive and planar cross-stratified beds represent a more proximal mouth bar setting (Orton & Reading 1993; Mellere et al 2002), and channelized units that occur at the top of the succession are interpreted as river distributary channels in a delta plain. This facies association is interpreted to represent a fluvio-dominated deltaic system or river delta (sensu Nemec 1990), fed by the braided channels of FA 2. This deltaic system was entering a shallow-marine setting open to the NE that most probably had a gentle gradient. This last interpretation is supported by the apparent absence of slumps and/or slides as well as the lack of evidence for Gilbert-type foresets (cf. Postma 1990). FA 4: distributary tidal channel. This facies association comprises fine- to coarse-grained sandstones in lenticular units that extend laterally for tens to a few hundreds of metres (Table 1). These sand-rich units show concave and erosive bases, and an upwards-fining and -thinning

trend. Planar-tangential cross-bedding are the dominant sedimentary structures, but sigmoidal and trough cross-bedding are also present (Fig. 6; e.g. RH and LR logs). Mud clasts and mud drapes occur on the foresets (Fig. 7e), and beds locally have rippled top surfaces. Convolute bedding is also common. Palaeocurrents are unimodal and towards the north (Fig. 8). Sandrich units are capped by bioclastic carbonates or densely bioturbated beds with Skolithos burrows. These sand-rich, lenticular units are frequently interbedded with finer-grained sandstones and siltstones with abundant ripple marks, rare Cruziana ichnofacies trace fossils and common mudcracks (Fig. 6, log RH). Lenticular units with concave and erosive bases are interpreted to represent channel-fill sequences. Two-dimensional transverse dunes, migrating dominantly northwards (basinwards) filled the channels. Regular, episodic changes in the hydrodynamic regime allowed ripples and mud drapes to form (Allen 1980) and therefore a tidal influence can be inferred (Shanley et al. 1992). However, the relatively small amount of mudstone suggests that flow velocity was high, and commonly precluded the accumulation or preservation of muds (Willis et al. 1999). Upwards-fining and -thinning trends occurred due to progressively local abandonment of the channel. These lenticular units are interpreted to represent ebb tidal channel-fill sequences in a subtidal setting. Although these sequences can be deposited as inlet tidal deltas (Sha & de Boer 1991), a distributary tidal channel origin in the active portion of a tide-influenced delta plain is favoured here (cf. Dalrymple 1992). This interpretation is supported by a close vertical relationship with tideinfluenced prodelta-delta-front facies of FA 5 (Fig. 6; e.g. log LR). Abandonment of the active channel led to accumulation of finegrained deposits in low-energy, subtidal-intertidal parts of the system. Towards the NW, tidal channel-tidal plain successions alternate with fluvial intervals, to produce a 120 m-thick succession (Fig. 6; log RH). This suggests that individual tidal channel-fill sequences were not deposited within incised valleys. FA 5: tide-influenced prodelta-delta front. Facies association 5 has only been recorded in Loma Rayoso section (Fig. 6; log LR, Table 1). This FA is composed of upwards-coarsening units that attain a maximum thickness of 12 m. At the base these units are composed of dark olive mudstones and parallel-laminated siltstones. Plant and comminuted organic debris are abundant and are typically associated with

ANATOMY OF A LOWSTAND SEQUENCE SET

diagenetic calcite nodules. Microfossils are absent. Dark mudstones are overlain by centimetre-thick, very-fine-grained sandstones interfingered with mudstones. Sandstones are predominately cross-laminated, with wavylenticular discontinuous beds with symmetrical and asymmetrical ripples. Cross-laminae dip predominantly to the north, but many have an opposite direction. These interbeds contain scarce undetermined shell fragments, rare Cruziana ichnofacies burrows and scattered synaeresis cracks. Large-scale inclined surfaces with convex-up geometry form the upper part of the thickest upwards-coarsening successions (Fig. 7f). Internal structures ranges from decimetre-thick cross-stratification with mud drapes in the upcurrent portion that pass to previously described interbeds in the downcurrent direction. Internal cross-strata dip in the same direction as the large-scale inclined surfaces (towards the north, Fig. 8), but locally, small-scale cross-stratified beds show evidence of flow reversal. Abundant plant material, synaeresis cracks, the trace fossils association and the scarcity of invertebrate fossils indicate that this facies association was deposited in a stressed marine environment which experienced salinity fluctuations (Bhattacharya & Walker 1991; Ftirsich 1995). Fluvial influence was present within this depositional setting, but was less important than in FA 3, allowing the development of mixed tide/wave and tidal deposits in a progradational deltaic system (cf. Dalrymple 1992). Dark mudstones accumulated from suspension in a steady, quiet prodelta environment, not suitable to support large benthic communities. Laminated siltstones and interbedded rippled sandstones and mudstones result from rapid, frequent alternations in current strength (Nio & Yang 1991). North-dipping cross-lamination in discontinuous sandstones record sediment transport during the dominant ebb tides, while mud drapes settled from suspension when currents waned. These heterolithic intervals record a distal-proximal delta-front setting, in a tide-influenced deltaic depositional system (cf. Dalrymple 1992; Willis et al 1999). Similar facies successions have recently been reported from delta-front settings of modern tide-dominated and tide-influenced deltas (Hori et al. 2002; Ta et al. 2002). Large-scale inclined deposits are interpreted as being deposited by subtidal migrating dunes (Allen 1980; Willis et al. 1999). Willis & Gabel (2001) have extensively discussed the formation of this kind of subtidal deposits, concluding that these macroforms are most likely to form in the upper parts

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of tide-influenced delta fronts, with episodic sediment supply. Estuarine facies associations FA 6: meandering fluvial plain-inner estuary. This facies association includes lenticular and tabular units that are generally much finer grained than those in FA 2 (Table 1). Lenticular units usually exhibit large-scale inclined strata (sensu Bridge 1993) typically dipping less than 15°. These elements are composed of trough cross-bedded strata that pass upward into ripple cross-laminated beds. Some units show a well-defined upwards-fining trend and often grade upwards into tabular deposits. Mud clasts are commonly found parallel to the inclined surfaces and, in a few cases, these surfaces are draped by mudstones that can be traced from the upper portion of the inclined surface into the basal part of the lenticular unit (Fig. 9a). The tabular packages consist of decimetre-scale heterolithic (lenticular and wavy) intervals, and massive or planar-laminated sandstone beds separated by mud drapes. Discontinuous sandstone lenses contain oscillatory and current ripples. Bidirectional ripples are frequent. Mud clasts generally occur at the base of beds and desiccation cracks are rare. A few Palaeophycus and Teichichnus burrows and gastropods were observed. These deposits are overlain by a facies association that represents central-outer estuarine conditions (FA 7, Fig. 6). The presence of large-scale inclined strata, concave bases and well-defined upwards-fining trends indicate that lenticular units were deposited within point bars in the convex margin of sinuous fluvial channels (Miall 1985; Bridge 1993), which flowed across a low-relief fluvial plain. The rare channelized deposits with continuous mud drapes from top to base may have formed within tidally influenced rivers (Thomas et al. 1987; Allen 1991). As these deposits typically pass upwards and down-dip into estuarine deposits these channels are interpreted to represent the upper reaches (sensu Allen 1991) or inner portion (sensu Dalrymple et al. 1992) of an estuarine depositional system. Vertically related interbedded sandmud couplets of tabular non-channelized units represent traction-suspension cycles, either in mixed tidal flats with dominant subcritical flow regime (ripples), or in sand flats under upper flow regime (plane beds). This interpretation is supported by the trace fossil associations (cf. Pemberton et al. 1992) and scarcity of macrofossils (cf. Ftirsich 1995) that suggest

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an impoverished fauna in a stressed environment, common in tidal flats. FA 7: central and outer estuary and transgressive coast. This association comprises a

variety of lithologies, from mudstones to poorly sorted sandstones, and more rarely skeletal calcareous sandstones (Table 1). Units exhibit either upwards-fining or, more frequently, upwards-coarsening trends. The latter consist of

Fig. 9. Main characteristics of estuarine and coastal plain facies associations, (a) Fluvial sandbodies with large-scale inclined strata of FA 6. In some examples surfaces are marked by mud drapes (arrows) from top to base of the body that may suggest tidal influence, (b) Upwards-coarsening successions interpreted to be accumulated in a tide-dominated estuary (FA 7). Marine ostracods were recorded for the basal grey mudstones, which constitutes the first occurrence in the locality. Upward increase in sand trend probably represents progradation of estuarine tidal bars. (Person for scale.) (c) Upwards-coarsening shallow marginal marine deposits capped by poorly exposed, thick shelf mudstones with ammonite horizons (dotted line). They are interpreted to represent outer estuarine settings and shore zones developed during transgression. (Person for scale.) (d) Heterolithic channels intercalated with dark grey massive mudstones that accumulated in low-energy coastal plains (FA 8) and embayments (FA9), respectively. Note the sharp base (arrows) associated with sand-rich bedsets; heterolithic beds become more abundant to the left and top. (e) Mud-draped wave ripples that are found filling heterolithic channels (FA 8) or forming the upper part of bay successions (FA 9). (Lens cap diameter is 5 cm.)

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mudstone intervals (Fig. 9b) with scarce marine ostracods at the base (Ballent pers. comm.). Mudstone intervals grade to sandstone-mudstone couplets with wave and current-ripple lamination that may pass upwards to tangentially cross-stratified packages of sandstones and bioclastic calcareous sandstones (Fig. 9b). Deposits of the FA 7 were recorded in the Southern and Central regions (Fig. 6; e.g. TR log and PN log). However, in the Central Region (downdepositional dip), sediments of this FA include a higher bioclastic content and a more diverse ichnofossil association (Cruziana and Skolithos ichnofacies), and are capped by a regional flooding surface (Fig. 9c). Cross-bedded deposits in this region suggest a main palaeoflow pattern towards the south (landward, Fig. 8). Thick mudstone intervals with marine ostracods indicate a muddy low-energy setting, with a marine connection. Interpretation of the accurate depositional environment (e.g. lagoon, bay or central estuary) requires a regional analysis (Yoshida 2000). Within this study, the spatial facies relationships with FA 6 indicate that this facies represents deposition in the central part of an estuary (sensu Dairymple et al. 1992). Sandstone-mudstone couplets grading upwards to tabular cross-stratified deposits are interpreted as distal and proximal parts of tidal estuarine sand bars, respectively (Allen 1991). In the Central Region, deposits of this facies association suggest a more marine influence and may represent outer estuarine facies (with some bedforms migrating landwards?), but also transgressive coastlines located in between the estuary mouths (cf. Allen & Posamentier 1993). According to estuarine models (e.g. Dalrymple et al. 1992; Reinson 1992), deposition of this facies association more probably occurred in a narrow, partially closed estuarine setting that was subjected to relatively moderate-high tidal influence and low wave influence. Low-energy coastal plain and embayment facies associations FA 8: low-energy coastal plain. Heterolithic lenticular units with concave-up erosive bases and transitional tops prevail in this facies association (Table 1, Fig. 9d). They exhibit 'largescale inclined heterolithic strata' (after Thomas et al. 1987; Bridge 1993), with sandier intervals more abundant near the concave margin of the bedsets, and sandstones and mudstones with flaser-lenticular bedding more abundant towards the opposite margin (Fig. 9d). Most sandstones beds show mud-draped ripples

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or low-angle cross-stratification. Trough-cross bedding is rare, and only occurs at the base of a unit, together with the coarsest sand and occasional ferruginous gravels. Planolites and Teichichnus burrows are common, as well as undetermined small shell fragments. The lenticular units are intercalated with massive dark grey mudstones that have abundant palynomorphs and exhibit no clear internal trend in grain size (Fig. 6; TRlog, 118-150m). The large-scale inclined heterolithic strata within channel-like bodies are interpreted as lateral accretion deposits within tidally influenced channels or creeks (Thomas et al. 1987; Allen 1991). The overall upwards-fining succession records the migration and gradual filling of channels that were characterized by low water discharge and a predominantly suspended sediment load (Miall 1996). Burrows and shell fragments indicate a proximity to marine-related water bodies. These tidally influenced channels developed in a low-energy coastal plain, most probably draining a lacustrine-palustrine setting (Haszeldine 1989) that supported abundant vegetation. FA 9: embayment/lagoon. This facies association is dominated by dark grey, carbonaceous shales and siltstones, with minor very finegrained sandstones. Units are up to 10 m in thickness and usually show an upwards-coarsening trend. Shales and siltstones are massive to weakly bioturbated, and contain a low-diversity and low-density marine ostracod association (Ballent pers. comm.) (Table 1). Beds with a high concentration of non-turriteliform gastropods are common. Sideritic nodules and abundant vascular plant fragments are frequent throughout. Sandstone beds are thin and have horizontal-wavy lamination, with oscillation ripples (Fig. 9e). They lack fauna but contain rare horizontal burrows (Table 1). This facies association occurs in the Southern Region and is stratigraphically related to FA 8 (Fig. 6; log TR, 110-150m). Sediments of this facies association mainly settled from suspension in a very low-energy environment. Abundant plant fragments, a scarcity of bioturbation, and low diversity and frequency of marine fossils suggest a marginalmarine environment, probably with brackish waters (cf. Pemberton et al. 1992) and a reducing subsurface environment. In this setting, wave energy was low and biogenic reworking was minimal. As mentioned above, interpretation of low-energy marginal marine settings strongly depends on regional and sequence stratigraphic context (Yoshida 2000). In this case, a wide

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embayment, lagoon or shallow (<10m deep) water body is envisaged, which constituted the down-dip culmination of suspended-load rivers of FA 8. Bay head delta progradation was probably responsible for the development of upwards-coarsening successions.

Shoreface - shelf fades associations FA 10: upper-middle shoreface. This facies association includes massive and cross-stratified beds (Table 1). Massive deposits consist of quartz-rich sandstones that contain few to abundant vertical burrows of Skolithos ichnofacies. Flat troughs oriented to the west (Fig. 8) with underlying shell lags were occasionally observed within these deposits. Cross-stratified, sandstone and calcareous sandstone beds show mostly tangential cross-bedding with abundant mud drapes (Fig. lOa), but trough cross-bedding was also observed. These beds usually pass laterally into heterolithic wavy intervals and may represent neap-spring tidal cycles (cf. Visser 1980), (Fig. lOa). Ichnofossils are rare to common in the sandstones (Skolithos ichnofacies) and rare in the carbonates (Thalassinoides). Palaeocurrents can be either unidirectional or bidirectional in each unit with a prevailing direction towards the north (Fig. 8). Vertical trends show an upwards transition from trough cross-bedding at the base, tangential cross-stratification in the middle with the coarsest, bioturbated sandstones occurring towards the top (Fig. 6; PC log, e.g. 278-290 m). This facies association represents deposition in a middle shoreface (subtidal)- upper shoreface environment. Tangential cross-bedded sands formed by the migration of two-dimensional (2D) sand dunes, mainly transverse to the dominant flow (towards the north). The presence of abundant mud drapes, thick-thin sandstones bundles and bidirectional flows attest a tidal influence in these deposits (Nio & Yang 1991; Shanley et al. 1992). Marine fossils and ichnofauna suggest an open-marine setting (Pemberton et al 1992). Tabular, usually erosive, units indicate that these mixed clasticcarbonate deposits accumulated first in subtidal channels (trough cross-beds) that vertically evolved into sandwaves (tangential cross-beds). Over time, bioturbated coarser deposits, with abundant vertical burrows formed in a waveand storm-influenced upper shoreface setting (Reading & Collinson 1996). Flat troughs with underlying shell lags in this setting are interpreted to represent the development and fill of short-lived rip channels during storm conditions.

FA 11: lower shoreface. This facies association is composed of very-fine- to fine-grained sandstones and oolithic calcareous sandstones (Table 1). Units of this FA exhibit either a preponderance of hummocky cross-stratification (HCS), or a high proportion of ripple cross-lamination. In HCS-dominated units, sandstone beds (10-60 cm thick) are composed of symmetrical hummocks (Fig. lOb) with minor internal, convex-up surfaces. Beds show basal erosive surfaces marked by bioclastic material and/or mud rip-up clasts. In ripple-dominated units, sets are less than 2 cm thick and have flat-undulatory basal surfaces. Cross-laminae dip in opposite directions in adjacent sets, and mud chips and drapes are preserved locally between sets (Fig. lOc). Few resting traces were observed in this facies association, although there is a high density of trace fossils (Fig. lOc), predominantly belonging to the Skolithos ichnofacies with Cruziana also common. HCS-dominated units commonly show an upwards-thickening trend and are almost exclusive to the NE part of the study area (BL and CP sections in Fig. 5). Ripple-dominated units are more common towards the Central Region. HCS-dominated units are interpreted to have been deposited above fair-weather base, in a storm-dominated lower-shoreface environment (Walker & Plint 1992). Sandstone units with a high proportion of ripple cross-laminations were also deposited in a shoreface setting, but related to a shore zone where frequency and magnitude of storms were not as important as in the first case. Ripple cross-lamination represents the migration of weakly asymmetrical ripples produced by combined oscillatory and unidirectional flows (cf. Myrow & Southard 1996). Although oscillatory flows suggest the persistence of storm waves in this environment, the common bidirectionality of cross-stratification indicates that deposition of these rippledominated units was significantly affected by reversing weak tidal currents. Accordingly, a low-energy, mixed (tide/wave) shoreface setting is envisaged for these deposits, similar to examples described by Prave et al (1996) and Molgat & Arnott (2001). FA 12: inner shelf. This facies association consists of thin interbedding of siltstones and veryfine-grained sandstones (Table 1). Micro-HCS and combined flow-ripples are dominant (Fig. lOd), but current ripples are also present (with northward-directed palaeocurrents). The sandstone/mudstone ratio varies between 30 and 70%. Few-abundant body fossils were observed (mainly bivalves), and trace fossils

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including Cruziana (residents) and Skolithos (opportunists) ichnofacies are very common, The facies association occurs in upwardscoarsening and -thickening packages and gradationally pass to FA 11.

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This facies association resulted from deposition on a storm-influenced/dominated shelf, between storm and fair-weather wave bases (Walker & Flint 1992). Hummocky cross-bedded and rippled sandstone beds record

Fig. 10. Main characteristics of marine facies associations, (a) Cross-stratified sandstones showing sand-rich parts (S) with mud drapes, which pass laterally into heterolithic (H) wavy intervals (FA 10). They resemble neap-spring tidal cycles and characterize middle (or subtidal) shoreface setting. (Hammer length is 33 cm.) (b) Planar-con vex-up (arrow), low-angle stratification that typifies large-scale hummocky cross-bedding of storm-dominated shoreface setting (FA 11). (Lens cap diameter is 5 cm.) (c) Small-scale cross-stratified sets of low-energy, tide/wave-influenced lower shoreface (FA 11). Cross-laminae (C) dip in opposite directions in adjacent sets, mud chips (M) are locally preserved and vertical burrows (B) are common. (Pen length is 14 cm.) (d) Sandstones and mudstones of the inner shelf facies association (FA 12). Small-scale hummocks with either planar (grey arrows) or irregular bases (white arrows) are very common. (Hammer length is 33 cm.) (e) Plan view of a shell bed with high degree of articulation (e.g. trigonids, T) and incrustation (e.g. cemented oysters) that intercalates among outer shelf mudstones (FA 13). Mixing of benthic fauna (T) and neritic fauna (ammonites, A) can be found. (Hammer width is 3 cm.) (f) Tabular and massive oyster accumulations typical of the carbonate interval of the Mulichinco Formation (FA 14). Inset shows densely-packed oysters in a micritic matrix, where some shells are articulated (arrows). (Coin diameter is 2.3 cm.)

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storm events while siltstones accumulated during fair weather. FA 13: outer shelf. This facies association is composed of green, massive mudstones and some marls, in thick units (Table 1). Sparse, usually articulated, macrofossils, foraminifers and ostracods are common. Within the mudstones, two different shell-bed types intercalate. The first type consists of thin, erosive, discontinuous beds with small bioclasts (< 1 cm long), usually con vex-up and disarticulated. They can grade upwards to siltstones with parallel lamination or current-combined flow ripples. The second shell-bed type is thicker and coarser grained (Table 1), with abundant micritic matrix, and with a high degree of articulation and incrustation. Fossil content is moderate-high with a mixing of infauna, epifauna (cemented oysters) and/or neritic fauna (ammonites) (Fig. lOe). Calcite nodules may be present at the base and glauconite grains have been observed in thin sections. This facies association is interpreted as representing deposition on a marine shelf below storm wave base. Thin bioclastic interbeds are considered as distal tempestites ('event shell beds', Kidwell I99la) deposited from low-density storm-induced currents with significant remobilization of fragments. Articulated-fossil-bearing beds are interpreted as parautochthonous shell concentrations that also accumulated below or near storm wave base (Kidwell 1991 b). However, in this case, reduced sediment supply produced faunal concentrations, taphonomic feedback (sensu Kidwell 199Ib) and authigenic minerals. Carbonate facies association FA 14: carbonate ramp with oyster accumulations. This facies association crops out in the Northern Region associated with siliciclastic shoreface-shelf-facies associations (Fig. 6; PC and LP logs). Upwards-coarsening cycles occur in the eastern part (PC log), but towards the west, cycles are less obvious (LP log). Coarsening-upwards examples encompass marls, bioturbated wackestones and densely packed oyster shell beds (boundstones) from base to top. Marls do not contain fossils except for some ammonites, wackestones have some thinshelled bivalves, and boundstones are composed of cemented oysters (± serpulids) that show pristine shell conditions (Fig. lOf). This facies has a tabular geometry and is laterally extensive over a few kilometres. Calcitic nodules are locally present at their bases. Occasionally, packstones

with reclined oysters, thick-shelled infaunal bivalves and echinoids can cap the cycles (Table 1). Massive beds are dominant. A radiolarian-foraminiferous microfossil association has been recognized throughout the beds (Ballent pers. comm.). These massive carbonate deposits record a long period of low clastic input. As mechanical structures are absent, the inferred carbonate depositional system is based on fossil associations and grain size (cf. Aberhan 1993). Macro- and microfaunal content attest to an open marine environment where abundant nutrients and good oxygenated conditions in the water column allowed radiolarian to flourish. Preserved benthic faunal associations are dominated by bivalves, and their distribution along the sea floor was controlled by the substrate and kinetic energy of the environment, with the most robust organisms (e.g. reclined oyster and thick-shelled bivalves) living in the coarsest and more mobile substrate (Aberhan 1993). Accordingly, upwards-coarsening cycles are interpreted to represent upwards-shallowing events in a low-energy, low-profile setting, where a weak wave regime was the only active hydraulic process. This depositional system is likely to resemble the epiric ramp model proposed by Lukasic et al. (2000) for cool-water interior seas. Sediment starvation and, probably, high stress marine conditions favoured low-diversity oyster accumulations over large areas. Sequence stratigraphy and palaeogeographic evolution The sequence stratigraphic analysis (Fig. 11) and palaeogeographic reconstructions (Fig. 12) of the Mulichinco Formation are based on the recognition of systematic changes in the sedimentary environments, the identification of key stratigraphic surfaces and interpreted changes in the patterns of accommodation distribution within the basin (cf. Van Wagoner et al. 1990). Accurate dating of systems tracts and key surfaces was possible through the refined ammonite biostratigraphy erected for the Valanginian of the basin (Fig. 4). The base of the Mulichinco Formation is marked by an unconformity formed in response to a relative fall in sea level during the Valanginian (Vergani et al. 1995). This drop in sea level was induced by tectonic inversion within the basin (Vergani et al. 1995). Deposits underlying the Mulichinco unconformity are characterized by inner- to outer-ramp sandstones and

Fig. 11. (a) North-south and (b) east-west correlation showing depositional environment distribution and sequence stratigraphic framework of the Mulichinco Formation. The datum coincides with the appearance of the first Pseudofravella angulatiformis ammonite horizon in the entire basin. It represents a maximum flooding surface and, according to Flint & Nummedal (2000), cannot be regarded as a horizontal surface. Inset: map of the study area showing the locations of cross-sections.

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Fig. 12. Schematic palaeogeographic evolution of the Mulichinco Formation.

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mudstones (Quintuco Formation) and basinal black shales (Vaca Muerta Formation, Fig. 2). The changes in depositional environments that occur across the unconformity indicate a sealevel fall of a least 100 m, with a comparable magnitude of sea-level rise occurring 2 Ma later, with the deposition of the deep-water shales of the overlying Agrio Formation (Fig. 2). Therefore, the Mulichinco is a package of shallow-marine and continental strata, encased within offshore deposits and is consequently interpreted as a lowstand wedge (sensu Van Wagoner et al 1988). The 2 Ma Mulichinco lowstand wedge (Figs 2 & 4) is part of a secondorder depositional sequence (sensu Vail et al. 1977). Different systems tracts (LST, TST and HST) were recognized within the Mulichinco lowstand wedge (Fig. 11). The Mulichinco lowstand wedge therefore comprises one complete third-order sequence (Fig. 11) and is termed a lowstand sequence set (sensu Mitchum & Van Wagoner 1991). Further subdivision into even higher frequency sequences is possible (Schwarz 2003) but is beyond the scope of this paper.

Sequence boundary In the Southern and Central regions the sequence boundary at the base of the Mulichinco Formation is represented by a dramatic shift in facies from outer ramp/basin deposits to alluvial deposits (Fig. 11 a). At least 100km (along strike) of the previous ramp-type depositional profile was subaerially exposed (Fig. 12a). No angular relationships were observed in outcrop, but truncation terminations are documented below the unconformity in seismic sections towards the south of the studied area (Vergani et al. 1995). In outcrop, the sequence boundary is represented by a low relief regional surface of erosion, overlain by a widespread fluvial sheet (Fig. 13a). Towards the northern Central and Northern regions there is no evidence of subaerial exposure associated with the sequence boundary (Fig. 12a). However, a sharp-based sandstone body directly overlies dark shales in the Puerta Curaco section (Figs 11 & 13b). At the base this package comprises hummocky cross-bedded sandstones with gutter casts, loading structures and abundant plant debris, and it is topped by a bioclastic lag followed by shelf mudstones (Fig. 6; PC log, 0-8 m). Deposition of this sandstone body was related to wave- and storm-influenced flows over a semi-consolidated substrate (Plint 1991), followed by rapid shoreface sedimentation in a low accommodation setting. Therefore, the surface at the base is

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interpreted as a regressive surface of marine erosion formed during forced regression (Plint & Nummedal 2000). According to Plint & Nummedal (2000), the last marine surface of erosion or the amalgamation of several surfaces (only one recognized in this study) should be taken as the correlative surface of the sequence boundary in the marine portion of the basin (Fig. 11). Further basinwards there is a gradual transition from dark shales to green mudstones and parautochthonous shell beds (e.g. CP log in Fig. 6). This transition represents the correlative conformity of the sequence boundary (sensu Van Wagoner et al. 1988) (Fig. 11). Ammonite horizons belonging to the Lissonia riveroi Zone were found both below and above the marine correlative surfaces of the sequence boundary. This high biostratigraphic resolution indicates that the time encompassed between the top of the previous second-order highstand systems tract (HST) and the base of the Mulichinco lowstand wedge was less than 0.5 Ma (Fig. 14).

Lowstand systems tract (LST) The LST is deposited in the time between the lowest relative sea level and the onset of rapid sea-level rise (Hunt & Tucker 1995) (Fig. 14). Basin floor fan or slope-fan deposits are usually absent in ramp-type settings (Van Wagoner et al. 1988; Posamentier & Allen 1993). This is the case for the Mulichinco Formation (Figs 11 & 12b), where the low relief profile of the basin precluded deep-marine fan deposition (Legarreta & Uliana 1991). Within the Mulichinco lowstand sequence set it is possible to recognize two different stacking patterns during the LST, informally called the Early and Late lowstand systems tracts (Figs 11 & 12). In the Southern Region, the Early LST is represented by a widespread sheet composed of highly amalgamated fluvial channel-fill sequences (FAs 1 and 2), between 20 and 35 m thick (Figs lla & 13a). This fluvial sheet extends laterally along strike for more than 30 km, and the proportion of gravelly deposits decrease in a down-dip direction (Fig. 12b). These fluvial deposits developed immediately after the onset of sea-level rise (Early LST, Fig. 14), when accommodation creation resumed in the alluvial region and slight alluvial aggradation took place. The transition from Early to Late lowstand conditions in this area (Fig. 11) is marked by an increase in the proportion of fine-grained, non-channelized deposits in the alluvial facies associations (Fig. 13a). This change suggests a

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Fig. 13. (a) Low-relief sequence boundary (SB) of the Mulichinco lowstand wedge in the Southern Region (view to the east). It represents an abrupt vertical transition from ramp marine deposits (Quintuco Formation) to coarse-grained fluvial strata. Note more amalgamated (Early) to less amalgamated (Late) fluvial lowstand deposits, (b) Sharp-based shoreface sandstone body at the base of Mulichinco Formation, covered by shelf mudstones (Northern Region, view to the east). It is bounded by the regressive surface of marine erosion (RSE) at the base, correlative with the sequence boundary (SB), (c) Transgressive-highstand stratigraphy in the Central Region (Pichi Neuquen locality, view to the south). Backstepping of facies associations during TST comprise distal fluvial (FA 6) to estuarine (FA 7) to outer shelf deposits (FA 13) (other legend as in Fig. 11). (d) Parasequence set with upwards-thinning and -coarsening trend in the Central Region (Loma Rayoso locality, view to the SE). They represent the progradational HST of the Mulichinco lowstand sequence set and are capped by the 'Master Transgressive Surface' (MTS) or basal surface of the long-term, transgressive event.

higher rate of alluvial-plain aggradation that can be related to an increase in the rate at which space to accommodate sediment is created (Marriott 1999) (Fig. 14). The Late LST is also characterized by a decrease in fluvial discharge rates, as early lowstand gravelly proximal braidplains (FA 1) evolved to more regionally distributed sandy distal braidplains (FA 2) (compare Fig. 12b and 12c). In the Central Region, the basal part of the Early LST comprises thin sandy alluvial facies (FA 2). These are overlain by a progradational deltaic succession that passes upward from fluvio-dominated prodelta to delta-front and delta-plain deposits (FAS) (Fig. lla). This deltaic system was fed by early lowstand

braided rivers (FA 2) that were related to the steepest fluvial gradients and the coarsest clastic input of all the sequence. Accordingly, the development of river deltas in the SW region (Fig. 12b) indicates that, although marine processes (tides, waves) were probably present during the early lowstand, high fluvial discharge was still the dominant process. The transition from the Early to the Late LST in the Central Region is marked by a change in the depositional stacking pattern and dominant processes at the shore zone. The earlier progradational trend passes upward to an aggradational stacking pattern (Fig. 11) where sandy distal braidplain sediments (FA 2) interbed with tidal delta-plain deposits (FA 4) and tide-influenced

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Fig. 14. Relative sea-level curve interpreted for the Neuquen Basin during the Early-Late Valanginian based on the sequence stratigraphic analysis conducted in the Mulichinco deposits. The figure also shows the position of systems tracts identified in the Mulichinco lowstand sequence set and their very well constrained biostratigraphic control. Timing according to the biostratigraphic scheme shown in Figure 4 (and legend as in Fig. 11).

prodelta to delta-front deposits (FA 5). Facies associations indicate the enhancement of tidal action at the shore zone, associated with the decrease in fluvial activity (e.g. fluvial discharge) that occurred during the Late LST. This, in addition to the fact that greater amounts of sediment were stored in the fluvial plain, allowed the reworking of earlier fluvial-dominated delta deposits into tidal-channel and delta-front deposits (Fig. 12c). Consequently, this tide-dominated deltaic system was developed when accommodation equalled or slightly outpaced sediment supply (Fig. 14). In the Northern Region the lowstand systems tract is characterized by storm- and wavedominated marine deposits (FAs 11-13), which exhibit the development of two parasequence

sets separated by a flooding surface (Fig. 11). The first parasequence set suggests a change from previous forced regression during relative sea-level fall, to normal regression during lowstand conditions (Flint & Nummedal 2000). The second parasequence set shows a greater proportion of shelf v. shoreface deposits that may suggest a less progradational pattern and is therefore correlated with the Late LST recognized towards the SW (Fig. 11). Overall, a non-deltaic linear coast is proposed for the lowstand system tract in the Northern Region (Fig. 12b, c), where storm and wave processes prevailed. These palaeogeographic conditions may have been produced by a combination of factors. The north-south shoreline orientation was probably more prone to the effect of the

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inferred winter winds coming from the west (Schwarz 2003). Moreover, storm processes were important during the LST in the Northern Region because inferred fluvial systems coming from the east were probably less developed and consequently exerted less influence on the shore zone than those developed in the SW. The ammonite record in the Northern Region (Fig. 11) indicates that the LST spans between the latest portion of the Lissonia riveroi zone up to the middle part of the Olcostephanus (O.) atherstoni Subzone (Fig. 14). Transgressive systems tract (TST) The transgressive surface (TS) is the base of the TST and is represented by the transition from an aggradational to a retrogradational stacking pattern as the rate of relative sea-level rise increases rapidly (Hunt & Tucker 1995) (Fig. 14). The TST is bounded at its top by the maximum flooding surface (sensu Van Wagoner et al 1988). It was not possible to pick a single surface in the study area and consequently a maximum flooding zone (MFZ) that contains this surface was identified (Fig. 11). Within the Mulichinco lowstand wedge, a backsteeping of facies associations is clearly recognized in the Central Region (Figs lla & 13c). Previous late lowstand braidplains (FA 2) and tidal delta plains (FA 4) evolved to meandering fluvial plains (FA 6), and ultimately to estuarine systems (FA 7) prior to flooding and sudden development of offshore conditions (FA 13) across the area (Figs 11 & 12d). Estuaries are interpreted here to be mainly narrow and elongated perpendicular to the shoreline (Fig. 12d). The MFZ is represented in this region by shelf deposits (Fig. 13c) that include thin limestones with abundant ammonites of the Karakaschiceras attenuatus Subzone (Schwarz 2003) (Fig. 4). This ammonite marker allows accurate dating of the TST between the later part of the Olcostephanus (O.) atherstoni Subzone and the basal part of the Karakaschiceras attenuatus Subzone (Fig. 14). In the Southern Region, the boundary between the LST and TST is defined by a rapid change in the fluvial style and in the proportion of floodplain deposits (Figs 6 & lla). Fluvial systems of the TST (FA 6, Fig. 12d) were more sinuous and had a lower sediment transport capacity than the lowstand rivers (FA 2). This is interpreted to be a consequence of reduction in fluvial gradient (cf. Shanley & McCabe 1994). Channelized deposits of the TST also had a lower interconnection than the sandbodies of the preceding LST due to a higher aggradation

rate. Central estuarine deposits (FA 7) developed on top of fluvial—inner estuarine deposits (FA 6) record a relative deepening of the system in the Southern Region as transgression continued (Figs 11 & 12e). However, the lack of a clear change to progradational patterns in this region (see below) makes the identification of the MFZ more difficult here. In the Northern Region, the flooding surface that occurs at the top of the second parasequence set is interpreted to represent the TS within the marine environment (Fig. 11). The TST consists of marls and parautochthonous limestones with glauconitic aggregates (FA 13) at the base, followed by carbonate cycles (FA 14) with abundant fossil invertebrates (Figs 11 & 12d, e). These deposits accumulated in a low-energy ramp setting subjected to very low clastic input. Overlying mixed clastic-carbonate cycles (Fig. 6; log PC, 170-180m) are interpreted as the transition to highstand conditions and, hence, represent the MFZ in this region (Fig. 11). Highstand systems tract (HST) The HST is a progradational set of deposits laid down during a progressive reduction in accommodation after the maximum flooding zone or surface (Hunt & Tucker 1995) (Fig. 14). In the eastern part of Central and Northern regions, HST deposits are easy to identify (Fig. 11). These deposits consist of several parasequences that show a general upwards-coarsening and -thinning trend with a higher proportion of upper shoreface deposits towards the top (Figs 11 and 13d). This internal arrangement suggests an early aggradational-late progradational stacking pattern, but progradation was insufficient to develop continental facies in the study area (Fig. 12f). Parasequences are mainly composed of storm- and wave-dominated marine sediments (FAs 10-13), with some influence of tides in the middle shoreface (FA 10). Deposits are composed of both sandstones and calcareous sandstones, which suggest a limited sediment supply during the HST (at least lower than during the LST). In the western part of the Central and Northern regions, coeval sediments of previously described HST deposits are characterized by thick outer-shelf deposits (mainly FA 13, Fig. lib). Accordingly, a clear proximal (east)-distal (west) marine trend is recognized during the HST of the Mulichinco lowstand sequence set (Fig. 12f). In the Southern Region, identification of unambiguous HST deposits is problematic. During the last stage of the Mulichinco lowstand

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wedge this region evolved into a low-energy coastal plain, with tidally influenced meandering channels (FA 8) entering wide shallow bays or embayments (FA 9, Fig. 12f). Within HSTs, coastal-plain depositional systems are usually present as progradational delta-top successions (Posamentier & Vail 1988). However, in the Mulichinco sequence the depositional system is constantly retrograding and becoming more marine influenced upwards (bay deposits overlie coastal plain deposits, Fig. 11 a). This suggests that creation of accommodation outpaced the sediment supply in the Southern Region during the last stage of the Mulichinco lowstand wedge, and HST deposits were not formed (Fig. 11). The upper boundary The Mulichinco lowstand sequence set forms the lowstand wedge of a second-order sequence. The exact position of the upper boundary of this lowstand wedge, and therefore the transition to the second-order TST, is unclear, but a flooding surface near the top of the Mulichinco Formation is the most likely candidate. This surface is better seen in the central and NE part of the study area, where shelf deposits overlie shoreface deposits (Figs 11 & 13d). This flooding event marks the acceleration in the rate of long-term relative sea-level rise and therefore corresponds to a transgressive surface, called herein the Master Transgressive Surface (MTS, Figs 11 & 14). Towards the west the MTS is marked by the first occurrence of basinal dark shales and mudstones on top of outer-shelf green mudstones (Fig. lib). The MTS is correlated in the entire area as it is associated with the first occurrence of a new olcosthepanid ammonite marker in the basin (Figs 11 & 14). Consequently, the HST of the Mulichinco lowstand sequence set expands between the basal-middle Karakaschiceras attenuatus Subzone and the basal Olcostephanus (Viluceras) permolestus ammonite subzone (Fig. 14). As the transgression proceeded, dark shales (suggesting anoxic bottom conditions) accumulated almost throughout the entire Neuquen depocentre (Figs 2 & 11). This final palaeogeographic situation is associated with another ammonite horizon that marks the beginning of the Pseudofavrella angulatiformis zone (Figs 11 & 14). In the distal marine region of the study area (e.g. La Parva section, Fig. lib), a transitional relationship between dark shales of the two different biostratigraphic zones is recorded. Towards the marginal areas, however, a metrescale shallow-water carbonate bed occurred

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encased within these relatively deeper marine deposits (Figs 6 & 11). This unit is interpreted to record a minor relative sea-level drop or decrease in the rate of long-term sea-level rise (Fig. 14). Accordingly, the base of the carbonate bed represents a minor regressive surface of marine erosion, which pass down-dip to a correlative conformity (Fig. lib). Discussion: classic model v. the technically influenced Mulichinco lowstand wedge In the classic sequence stratigraphic model (Posamentier & Vail 1988), a lowstand wedge was originally defined as a regressive stratigraphic unit characterized by a progradational pattern, and bounded at the base by the sequence boundary and at the top by the transgressive surface (Posamentier & Vail 1988). In the model erected for intracratonic basins characterized by a ramp profile, the lowstand wedge was composed of one or more marine parasequence sets that correlated up-dip with fluvial-estuarine deposits developed within incised valleys (Posamentier & Vail 1988; Van Wagoner et al. 1988). Nevertheless, the same authors emphasized that the model was flexible enough to accommodate variations in key parametres such as subsidence and sediment supply (Posamentier & Vail 1988). The Mulichinco lowstand wedge is bounded at the base by a sequence unconformity and at the top by a master transgressive surface (Fig. 11), and was developed in a back-arc basin with a ramp-type physiography (Fig. 3). However, the evolution of the Mulichinco lowstand wedge occurred during a period of active tectonics. The accumulation of the unit took place immediately after a major uplift event and during faultcontrolled subsidence (Vergani et al. 1995) (Fig. 2). Tectonics can strongly affect key parametres such as subsidence and sediment supply (e.g. Frostick & Steel 1993), while basin physiography can control the amount of erosion produced during a sea-level fall (Posamentier & Allen 1993). Accordingly, the non-marinemarine deposits of the Mulichinco lowstand wedge provided an excellent opportunity to explore variability within lowstand wedge systems. Instead of being a single regressive stratigraphic unit, the Mulichinco lowstand wedge comprises different systems tracts characterized by particular stratigraphic patterns (Early and Late LST, TST and HST, Figs 11 & 12), which developed under specific conditions of

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accommodation and sediment supply (Fig. 14). The Mulichinco lowstand wedge is therefore better envisaged as a lowstand sequence set. Alluvial deposits of the lowstand wedge were not accumulated within incised valleys. Rapid uplift, together with the low-gradient physiography of the basin, might have precluded the development of deep incisions during sequence boundary generation (cf. Posamentier & Allen, 1993). When accommodation resumed in the alluvial region, an amalgamated fluvial sandstone sheet developed on top of a widespread low-relief sequence boundary. Despite this fact, and as proposed in the models, these deposits were in time covered by estuarine deposits (TST,Fig. 11). Shoreline sedimentation was strikingly variable in different areas of the Mulichinco lowstand wedge, particularly during the LST (Fig. 12b, c). In the SW shore zone, fluvio-dominated deltas were deposited during the Early LST, fed by high-energy fluvial briadplains (Fig. 12b). These environmental conditions, characterized by the steepest fluvial gradients and the coarsest clastic input of all the sequence, were probably driven by significant tectonically derived topography developed in the west (Fig. 12a,b). Conversely, towards the NE sector of the Mulichinco depositional area, the facies associations suggest lower energy, non-deltaic coasts with an apparent lack of fluvial influence during the LST (Fig. 12b, c). It is proposed here that this situation might have been associated with a lower (if any) topographic relief created during the tectonic inversion in the eastern side of the basin. General shoreline trajectory was not uniform during the latest stage of Mulichinco lowstand wedge evolution. On the NE side, an overall progradational stacking pattern, and therefore a clear HST, is recognized after maximum transgression (Figs 11 & 13c). In contrast, in the SW part of the study area, wide embayments with negligible fluvial input evolved from previous narrow estuaries during the late stage of Mulichinco accumulation (Fig. 12e, f). This palaeogeographic evolution suggests an increase in marine influence with time, which was possible due to continuous slow retrogradation of depositional systems after the maximum rate of transgression. The difference in stacking pattern and resultant shoreline trajectory in certain areas of the Mulichinco lowstand wedge is interpreted as a response of different rates of subsidence across the depocentre. Intrabasinal variations in sediment supply could have also generated the observed pattern, but clastic input is interpreted to have been similar (limited) all across the

depositional area during the last stage of evolution (Fig. 12f). Higher subsidence in the SW resulted in continued transgression, while lower subsidence in the NE area resulted in progradation as the rate of sea-level rise started to slow (Fig. 11). Greater subsidence in the west also produced a shifting of the depocentre from south to west through time (compare Fig. 12c and 12f). Increasing subsidence in the west during the accumulation of the Mulichinco lowstand sequence set may be attributed to fault-controlled subsidence (Fig. 12e, f). Mechanical relaxation of previously inverted fault systems (Fig. 12a) could be invoked as the mechanism responsible for the differential subsidence (cf. Vergani et al 1995). The Mulichinco lowstand wedge highlights the variability of depositional settings and stratal patterns that can be developed within tectonically influenced lowstand wedges accumulated in basins with ramp-type physiography. Tectonics and eustasy can act together or separately to produce relative changes in sea level, which in turn control accommodation (Vail et al. 1991). In the case of the Mulichinco Formation it is proposed that tectonic inversion may have heavily modified or even overprinted the eustatic signature within the Neuquen Basin during the Valanginian. Conclusions • The Valanginian Mulichinco Formation was deposited after a major relative sea-level fall, partially (if not totally) triggered by a tectonic inversion pulse. Fourteen facies associations were identified in the Mulichinco lowstand wedge. They represent accumulation ranging from gravelly fluvial braidplains to outer-shelf marine settings. Sequence stratigraphic analysis, supported by reliable biostratigraphic zonation, resulted in the identification of early and late lowstand, transgressive and highstand systems tracts. The Mulichinco lowstand wedge comprises one third-order sequence that lasts about 2 Ma and represents the lowstand sequence set of a second-order sequence. Spatial and temporal distribution of the sub environments in the Mulichinco lowstand wedge shows a high degree of variability. River deltas developed during Early LST in the SW shore zone. Storm-dominated, nondeltaic coastlines prevailed in the NE area during the LST and HST. Carbonate ramp and narrow estuarine settings formed during the TST. Wide embayments developed in

ANATOMY OF A LOWSTAND SEQUENCE SET

the south during the last stage of evolution. Evidence for tidal activity occured not only in the TST, but also in the late lowstand interval (tidal deltas) and the HST. Alluvial deposits of the Mulichinco lowstand wedge were not developed within incised valleys, but they accumulated on top of a low-relief sequence boundary. Tectonic ally derived topography, basin physiography and fault-controlled subsidence were the main controls on the evolution of the Mulichinco lowstand wedge. The Mulichinco Formation provides some insights about the variability of depositional environments and stacking patterns that can be expected within lowstand wedge systems.

This study is part of the work submitted by E. Schwarz as a PhD project, which was funded by a postgraduate fellowship from the CONICET. CONICET and the National University of La Plata also supported field work and E. Schwarz's fellowship in Liverpool. E. Schwarz sincerely thanks his supervisor L. Spalletti for his encouragement, and many fruitful comments and discussions. The authors also thank S. Ballent (National Museum of La Plata, Argentina) who carried out the microfossil study. All the people who collaborated with E. Schwarz during his PhD project are also thanked, especially G. Veiga. S. Flint and members of the Stratigraphy Group in Liverpool are also acknowledged for discussions. J.A. Howell's fieldwork in Argentina was supported by the Stratigraphy Group. J. Hechem and an anonymous reviewer provided many helpful suggestions that improved the final manuscript.

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Anatomy of a mixed marine-non-marine lowstand wedge in a ramp setting. The record of a Barremian-Aptian complex relative sea-level fall in the central Neuquen Basin, Argentina GONZALO D. VEIGA1"3, JOHN A. HOWELL2'4 & ANNA STROMBACK2'5 1 CONICET - Concejo Nacional de Investigaciones Cientificas y Tecnicas, Argentina 2 Stratigraphy Group, Department of Earth Sciences, University of Liverpool. 4 Brownlow Street, Liverpool L69 3GP, UK 3 Present address: Centro de Investigaciones Geologicas, Calle 1 #644, B1900TAC La Plata, Argentina (email: [email protected]) Present address: Centre for Integrated Petroleum Research, University of Bergen, Allegaten 41, Bergen, Norway 5 Present address: Statoil, Forushagen 4035, Stavanger, Norway Abstract: During the Cretaceous, western Argentina was occupied by the Neuquen Basin, a back-arc-foreland basin that was open through the proto-Andes to the Pacific Ocean in the west. The Neuquen Basin contains a thick succession of sediments that include the offshore marine deposits of the Agrio Formation. These deposits represent a time when the arc was an island chain and the Neuquen Basin was freely connected to the Pacific. This offshore marine succession is punctuated by two intervals of arid continental deposits that represent major, second-order, relative falls in sea level. In both of these cases there is no evidence of tectonic uplift or angular truncation along a basal bounding unconformity. The upper of the two lowstand wedges is characterized by a complex arragement of shallow-marine and continental deposits. Shallow-marine deposits sharply overlying offshore shales and capped by a master sequence boundary are interpreted as falling-stage deposits recording a complex relative sea-level fall. On top of a regional erosion surface, a drying-upwards succession of fluvial-aeolian deposits is developed, recording a fully non-marine stage in the evolution of the basin. These deposits are overlain by a marginal marine evaporite succession. The absence of a return to fully open-marine conditions is attributed to uplift in the Andes and marks the transition of the Neuquen Basin from a back-arc to a foreland system. This succession has important implications for the basin's evolution and in the timing of the uplift of the Andes, is a very spectacular example of a lowstand wedge and is also a major hydrocarbon reservoir.

The Lower Cretaceous of the Neuquen Basin includes a thick succession of deep-marine shales that are punctuated by several intervals of shallow-marine and continental deposits. These intervals have been interpreted as major lowstand wedges produced by significant relative sea-level falls (Legarreta 2002). The aims of this paper are to document the internal architecture and bounding surfaces of the uppermost of these wedges, a continental succession of the Huitrin Formation, which sharply overlies offshore marine deposits of the Agrio Formation and is overlain by evaporites. Lowstand intervals have been documented previously

from the Neuquen Basin, but there is controvers as to the connection of the basin to the Pacific during their deposition. Some authors have proposed a 'Messinian-type' model in which the basin was totally cut off (Mutti et al. 1994; Legarreta & Uliana 1999). Others have documented equivalent marginal marine deposits in the centre of the basin (in Mendoza Province, see Fig. 1), indicating at least a partial connectio to the open ocean (Veiga & Vergani 1993; Sagasti 2002). This distinction has very important implications for our understanding of rates and magnitudes of relative sea-level fall, and also for the timing of Andean uplift.

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 139-162. 0305-8719/05/$15.00 © The Geological Society of London 2005.

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Fig. 1. Map of the Neuquen Basin showing the location of the studied area.

The lowstand responsible for the Huitrin Formation differs from previous lowstands in the Lower Cretaceous of the Neuquen Basin (e.g. Kimmeridgian and Hauterivian) because deep-marine conditions are not re-established following deposition of the lowstand wedge. Instead, during the subsequent relative sea-level rise there was a restricted connection with the open-marine environment that permitted the accumulation of a thick evaporitic succession. This suggests that the mechanisms that drove this relative sea-level fall might have been different to the ones that produced the previous stages of sea-level lowering, suggesting changes in basin-scale factors that might have controlled the development of these deposits. Furthermore, a complex array of falling-stage deposits are present in the transition from

deep- to non-marine facies, suggesting a composite sea-level fall event rather than a simple, rapid sea-level drop.

Geological setting and stratigraphy The Neuquen Basin is located in west-central Argentina (Fig. 1) and is characterized by a complex tectono-stratigraphic history. The basin was initiated during the Late Triassic by the extensional collapse of a Late Palaeozoic orogen (Franzese & Spalletti 2001). The onset of subduction on the western margin of Gondwana resulted in the development of a magmatic arc. During most of its Mesozoic history, the basin existed as a back-arc basin, open to the Pacific through gaps in the volcanic arc (Spalletti et al 2000). The changing nature of this

SEDIMENTOLOGY OF A LOWSTAND WEDGE

connection is key to the development of the basin fill. Eustatic sea-level fluctuations and uplift in the arc periodically affected the connection between the basin and the Pacific (Legarreta & Uliana 1999). Uplift and inversion in the Andes during the Cretaceous finally isolated the basin from the marine realm and, with further compression, it evolved from a back-arc into a foreland basin. Continued compression and uplift produced the very tight, isoclinal folding that characterizes the outcrop belts observed today. The study interval is Barremian in age. During the preceding Hauterivian time deposition in the central part of the Neuquen Basin was characterized by the accumulation of the ammonitebearing black shales of the Agrio Formation (Fig. 2). These off shore-deep-marine depositional conditions were briefly punctuated by the development of a non-marine lowstand. This lowstand, termed the Avile Member, lies within the Agrio Formation, and is composed of up to 100 m of aeolian and arid fluvial deposits in a wetting-upwards (sensu Clemmensen et al. 1989) succession (Veiga et al. 2002). Overlying the Avile Member there is a sharp transition back to offshore black mudstones of the Upper Member of the Agrio Formation (Fig. 2) (Veiga et al. 2002). The base of the study interval occurs towards the top of the Agrio Formation

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where, locally, shallow-marine deposits sharply overlie the offshore deposits that characterize the majority of the formation. The sharp and locally erosional nature of this contact indicates that an important change in depositional conditions occurred prior to the deposition of the overlying Huitrin Formation. The Huitrin Formation unconformably overlies the Agrio Formation (Fig. 2) and has been subdivided into three members (Groeber 1946). The lowermost Chorreado Member is only locally developed, and consists of shallow-marine carbonates and evaporites. Overlying the Chorreado Member, or where the Chorreado Member is absent on top of the black shales of the Agrio Formation, is the Troncoso Member. The Troncoso has a wider distribution than the Chorreado Member and has been subdivided into two portions based on lithology and stratigraphic position. The lower portion is a non-marine siliciclastic unit (Troncoso Inferior) and the upper an evaporitic succession (Troncoso Superior). The siliciclastic deposits of the Troncoso Inferior show a more restricted distribution than the overlying evaporites that occur basinwide. In the central Neuquen Basin the Huitrin Formation ends with the carbonate deposits of the La Tosca Member. This study concentrates on those deposits that lie between the surface that

Fig. 2. Chronostratigraphic chart for the Lower Cretaceous of the central Neuquen Basin showing the distribution of the lithostratigraphic units involved in this study. H*; Hauterivian; A*, Albian. Modified from Legarreta & Gulisano (1989); age scale from Gradstein et al. (1995), age of units taken from Aguirre Urreta et al. (1999).

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marks the end of deep-marine conditions, either towards the top of the Upper Member of the Agrio Formation or where this shallow-marine interval is absent at the base of the Huitrin Formation, and the transgressive surface at the base of the Troncoso Superior (Fig. 2). Several units within this interval, and especially the Troncoso Inferior, are important hydrocarbon reservoirs in the subsurface (Comeron 1990; Veiga & Rossi 1992; Vergani et al 2001, 2002; Barrionuevo 2002). However, previous studies of the outcrops, especially the significance of the surfaces that bound the different lithostratigraphic units, have been sparse and contradictory (cf. Legarreta et al. 1993; Legarreta 2002). Descriptions of the Huitrin Formation have also been presented from the northern part of Mendoza Province, to the north of the study area (Legarreta 1985, 1986). The Chorreado Member was previously studied by Gutierrez Pleimling (1991). Previous studies (Legarreta & Gulisano 1989) placed the study interval in the final stage of the Upper Mendoza Mesosequence (Ms4 Sequence) and the lower portion of the Huitrin Mesosequence (HI Sequence). These authors place the boundary between these two mesosequences at the stage of maximum marine regression and tentatively correlate it to the 112 Ma third-order sequence boundary of Haq et al. (1988). Biostratigraphic studies (Aguirre-Urreta & Rawson 1997; Aguirre-Urreta et al 1999) documented ammonites of the Paraspiticeras groeberi Zone in the Upper Member of the Agrio Formation indicating an age as young as lower Barremian. Location and dataset The study area is located in the central part of the Neuquen Basin, in the northern part of the Neuquen Province in west-central Argentina (Fig. 3). This area is located in the Andean sector of the basin and is included in the Agrio Fold and Thrust Belt, where numerous Tertiary structures have deformed the Mesozoic fill of the basin into tight, N-S-oriented folds. The available outcrops cover a wide range of palaeogeographic settings, and make possible the characterization of the lateral and vertical variations in the sedimentary environments. Twenty detailed sedimentological sections were measured at 10 localities (Figs 3-5). In addition to detailed facies analysis of logged sections, architectural panels were made to define the two-dimensional (2D) geometries of the deposits. Key sequence stratigraphic surfaces were identified within the main outcrops and correlated between the sections. These correlations

were used to supplement the relatively poor biostratigraphy within the study interval. Facies associations and sedimentary environments Sedimentary facies were defined on the basis of textural properties and sedimentary structures. These facies are summarized in Table 1. The facies have been grouped into six facies associations defined on an interpretation of the depositional environment (sensu Reading 1986). The facies associations are detailed below. Open-marine facies association Bituminous black shales, carbonate shales and marls with minor wackstones and packstones, and occasional grainstones, boundstones and siltstones comprise the majority of the upper part of the Agrio Formation in the central part of the study area (Spalletti et al. 200\b). Ammonites are common, and marine reptiles and other open-marine fauna have also been documented. In the southern part of the basin, similar shales are interbedded with thin sandstones containing wave ripples and hummocky cross-stratification (HCS) in coarsening-upwards successions 1.514.5 m thick (Spalletti et al 200la). The facies within the central part of the basin are interpreted as offshore deposits, laid down below the storm wave base in an open-marine depositional environment. The deposits within the southern part of the basin are also interpreted as open-marine sediments, although the presence of wave ripples and HCS indicate slightly shallower water depths, above storm wave base. The upwards-coarsening successions are interpreted to represent the distal portions of shoreface successions, indicating that this locality was somewhat closer to the basin margin. The shallower marine part of this facies association occurs exclusively in the upper part of the Agrio Formation. Overall, the association represents marine conditions that prevailed prior to the relative fall in sea level at the base of the Huitrin Formation. Further details on this facies association can be found in Spalletti et al (2001^ b) Siliciclastic shallow-marine facies association This association is subdivided into two subassociations, both of which represent a different component of the shallow-marine system. All of the siliciclastic shallow-marine deposits have

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Fig. 3. Simplified geological map of the study area with the distribution of Mendoza Group and Huitrin Formation outcrops and studied localities. BHT, Balsa Huitrin; CHM, Chos Malal; CH-ME, Chacay Melehue; CH-MEN, Chacay Melehue North; CUR, Cerro Curaco; EAT, Tril Anticline East; ECR, Cerro Rayoso; ECR-S, Cerro Rayoso South; FMO, Filo Morado; TM, Tricao Malal; TMS, Tricao Malal South; TRL-E, Pampa de Tril East; TRL-S, Pampa de Tril South; TRL-W, Pampa de Tril West.

a limited lateral extent and are discontinuous, even between closely spaced outcrops. All examples overlie, with sharp and locally erosive contacts, the offshore, open-marine deposits of the Agrio Formation described previously. Siliciclastic shallow-marine deposits are common in the southern and eastern sectors of the study area. Where both sub-associations are present, the offshore transition-lower-shoreface deposits pass upward into upper-shoreface deposits. Both sub-associations have been identified sharply overlying offshore marine facies of the Upper Member of the Agrio Formation. Sub-association 1: offshore transition-lower shoref ace. Offshore transition-lower-shoreface deposits are restricted to the southern part of the studied area. They are present in the

southernmost localities (Balsa Huitrin and Cerro Rayoso), as well as locally in the Pampa de Tril area (Fig. 5). When present, this association is characterized by a 10-20 m-thick succession with a sharp-erosive basal surface. The sub-association is dominated by fine-grained bioclastic HCS sandstones in packages 10-40 cm thick. Bioclasts comprise disarticulated and fragmented bivalves and pentacrinites ossicles. Bioturbation is neither frequent nor intense, with Thalassinoides and Palaeophycus as the only ichnogenera identified. The HCS sandstones are interbedded with grey siltstones and mudstones up to 1 m thick. Wavy and lenticular bedded heterolithic intervals are also present in successions up to 3 m thick. Fine-grained sandstones with wave ripple and planar lamination in beds 1020 cm thick also occur, as do sandstones with

Table 1. Fades associations for the studied interval Facies association Open marine (OM)

Siliciclastic shallow marine (SSM)

Offshore transition (OT)/lower shoreface (LSF)

Upper shoreface (USF)

Carbonate shallow marine (CSM)

Fluvial

Lithology Black shales, carbonate shales and marls. Minor mudstones, wackstones and packstones. Subordinate siltstones, boundstones and grainstones Fine- to medium-grained bioclastic sandstones. Hetrolithic intervals. Siltstones and mudstones

Fine- to medium-grained sandstones. Thin intercalations of siltstones Grainstones, packstones and wackestones interbedded with fine-grained sandstones and siltstones. Boundstones

Coarse-grained braided (CGF)

Coarse- to fine-grained sandstones. Abundant rip-up clast

Facies

Thickness Not considered in this study

HCS and small-scale trough cross-bedding in sandstones. Wave-ripple and planar lamination. Wavy and lenticular heterolithics. Gutter casts. Scarce bioturbation. Softsediment deformation Trough cross bedding and horizontal stratification in sandstones. Waverippled fine-grained sandstones. Laminated siltstones Oolitic, oncolitic and peloidal carbonates with small amount of bioclastic fragments. Laminated algal boundstones and massive - laminated siltstones Trough cross-bedding and horizontal stratification in sandstones, related to

10-20 m thick. LSF may sharply overlie OM deposits without OT accumulation

Up to 10 m. Transitionally from LSF deposits in an overall coarsening upward (CU) succession 2-30 m thick

0 to over 80 m. Related to valley incision in the south

conglomerates Subordinated conglomerates and mudstones

Fine-grained ephemeral (FGF)

Red and green mudstones. Fine- to coarse-grained sandstones. Scarce rip-up clast conglomerates

Aeolian (A)

Fine- to medium-grained sandstones. Dark mudstones

Evaporites (E)

Anhydrite and halite with thin intercalations of sylvite

lenticular erosive bodies. Mudstones mainly massive with desiccation cracks and rootlets associated with finegrained sandstones with current ripples. Intercalation of sandstones with windripple lamination Up to 50 m. Only present in Massive to laminated theNW mudstones with rootlets and desiccation cracks. Tabular fine-grained sandstones with current ripples and small-scale trough cross bedding or massive. Lenticular sandstone bodies with large-scale inclined strata. Isolated largescale planar cross-bedded sets with wind-ripple lamination. Variable over short Small- and large-scale distances 0-35 m trough and planar cross bedding. Wind-ripple lamination. Grain-fall and grain-flow laminae. Horizontal and wavy lamination. Softsediment deformation (SSD) and dune topography preserved at the top Not considered in this study

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Fig. 4. Sedimentary logs of the NW sector of the study area. Distribution of facies and facies associations. No horizontal scale. See Figure 3 for location.

SEDIMENTOLOGY OF A LOWSTAND WEDGE

147

Fig. 5. Sedimentary logs of the SE sector of the study area. Distribution of facies and facies associations. No horizontal scale. See Figure 3 for location.

small-scale trough cross-bedding, although the latter are rare. The basal contact of the subassociation is always sharp and is locally erosive with rounded, partially connected, finegrained sandstone-filled gutter casts (sensu Myrow 1992) up to 40 cm deep and 20 c wide (Fig. 6a). Immediately above the bas surface there is commonly an amalgamated

HCS bed up to 2 m thick (e.g. at Cerro Rayoso, Fig. 6b). The deposits of this sub-association represent accumulation above storm wave base in an off shore-transition-lower-shoreface environment. The predominance of storm- and wavegenerated sedimentary structures within the sandstones indicates that the depositional water

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Fig. 6. Shallow-marine facies association, (a) Sandstone-filled gutter casts at the base of the shallow-marine facies association in the Balsa Huitrin locality. The hammer is 33 cm long, (b) Amalgamated HCS sandstones in the Cerro Rayoso area. Note the sharp transition between offshore shales and lower shoreface sandstones, (c) Convolute sandstones of the upper-shoreface sub-association in the Porton de Tril area. The hammer is 33 cm long, (d) Corrugated algal boundstones of the carbonate facies association. Chacay Melehue sector. The pen is 15 cm long.

depths were shallower than those that existed during the deposition of the underlying Agrio shales. Gutter casts at the base of this interval indicate an initial episode of scouring, potentially promoted by storm-related downwelling flows during a period of relative sea-level fall (Flint & Nummedal 2000). Also, the amalgamation of HCS sandstones directly on top of offshore shales with the offshore transition association missing indicates an episode of loss of accommodation, suggesting a relative sea-level lowering. Sub-association 2: upper-shoreface e facies. Deposits interpreted as being laid down within an upper-shoreface setting were only observed in the Porton de Tril section (Fig. 5). At this locality a 10 m-thick sandstone succession is present, transitionally overlying offshoreshoreface-transition deposits characterized by laminated mudstones, siltstones and some intercalations of HCS sandstones with abundant syndepositional deformation structures (Fig. 6c). The sandstone succession is characterized by wellsorted medium-grained sandstones in beds up to 1 m thick. Planar stratification and trough cross-

bedding dominate this sequence, but some small-scale intercalations of wave-rippled finegrained sandstone were also identified. The absence of bioclasts and icnofossils throughout the section is conspicuous. These deposits represent accumulation in a shallow-marine environment. The presence of planar and trough cross-stratified sandstones suggests that these deposits might have accumulated in the surf and breaker zone of a relatively high-energy coast (Olsen et al 1999). The abundant soft-sediment deformation observed is linked to rapid loading onto saturated fine-grained sediments, possibly promoted by an increase in the rate of sediment supply (Fitzsimmons & Johnson 2000). The absence of bioturbation or bioclastic fragments in these facies contrasts with the rest of the shallowmarine deposits observed in the study area and suggests a high-energy environment with conditions not suitable for the proliferation of organisms. This can also be related to the development of more brackish conditions promoted by an increase in freshwater supply from fluvial distributaries.

SEDIMENTOLOGY OF A LOWSTAND WEDGE

Carbonate shallow-marine fades association Carbonate shallow-marine deposits characterize the Chorreado Member in the studied area (Figs 2 & 4). These deposits are restricted to the central part of the basin and are completely absent in the SE sector. They show considerable variations in thickness, ranging from 2 m in the NW of the study area to 35 m in the east. These thickness changes are not typically associated with marked changes in facies, although the unit contains large-scale sigmodial geometries, dipping in a NW direction (Gutierrez Pleimling 1991). This association is characterized by oolitic grainstones, packstones and wackestones that are interbedded with sandstones and calcareous siltstones. Algal boundstones are common in the central part of the basin and occur towards the top of the section in the NW (Fig. 6d). These deposits have been comprehensively described by Gutierrez Pleimling (1991) who interpreted them as part of a carbonate ramp system surrounded by an oolitic complex. The ramp deepened and prograded towards the NW and the sigmodial geometries have been interpreted as clinoforms. The shallow-water algal boundstones are interpreted as tidal-flat deposits and their presence in the distal NW part of the basin indicates the shallowing of the marine environment and the development of a regional unconformity within these deposits (Gutierrez Pleimling 1991). The low degree of bioturbation and the small proportion of bioclastic fragments suggest a restricted environment, with low proliferation of organisms as for the siliciclastic shoreface facies association.

Fluvial facies association The fluvial deposits are divided into two subassociations, coarse-grained braided and finegrained ephemeral, based on the type of river systems that is interpreted to have deposited them. There is distinct spatial and stratigraphic control on the distribution of these different systems. Coarse-grained braided. These deposits are present in almost every studied locality, although the thickness is highly variable. In areas such as the Colorado River area (Fig. 4) the system may reach up to 80 m in thickness, whereas in the southern sector the association is thinnest, ranging from 0-15 m (Fig. 5). The base of the interval is typically very erosional and in the SE of the study area much of the thickness

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results from the filling of scours, 10-15 m deep and up to 200 m wide, cut into the black shales of the Agrio Formation. The deposits of this sub-association are characterized by well-sorted, medium- to coarse-grained sandstones in lenticular bodies up to 3 m thick with well-developed large-scale trough cross-bedding. The bodies show a concave and erosive base, and are stacked with a high degree of amalgamation. The palaeocurrents for these deposits are strongly unimodal with a main transport direction towards the west and NW. Locally, the base of the trough cross-sets is characterized by intraclastic conglomerates with clasts up to 15 cm in diameter. These intraclast conglomerates are associated with the most erosive portions of the channel bodies. In addition to the intraclasts, bouldergrade conglomerates with clasts from the underlying Chorreado Member and Agrio Formation are also present in the basal portions of the Troncoso Member (Fig. 7a). Within the lower portion of the interval, coarse-grained sandstones with horizontal lamination form units up to 2 m thick (Fig. 7b). These units have a higher width/thickness ratio (>15) than the sand bodies higher in the succession. Claystones and siltstones are not common in this association. Where present they occur as thin discontinuous beds that lie between channel sand bodies. The mudstones lack primary depositional structures, probably due to biogenic and pedogenic processes, and contain rootlets and sandstone-filled desiccation cracks. Horizontally bedded, fine-grained sandstones with crosslamination are also locally interbedded with the mudstones. The coarse-grained deposits also erosively overlie locally pale-green, fineto medium-grained sandstones that are bimodally sorted with horizontal to low-angle crossstratification showing pinstriped laminae typical of very-low-amplitude wind ripples. The facies of this sub-association were deposited in a braided fluvial setting, and reflect a very low accommodation rate. The low aggradation rate promoted the amalgamation of channel sandstone bodies, and favoured the erosion and reworking of overbank deposits. Individual channels would have been transient features and the uniformity of palaeocurrent directions indicates that the majority of deposition occurred within downstream accreting bars in low-sinuosity channels. Features within the rare mudstones indicate that overbank areas were occasionally humid enough to sustain vegetation, although the presence of desiccation cracks indicates that these areas dried out. The interpretation of a predominantly arid climate is supported by the

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G. D. VEIGACTAL.

Fig. 7. Fluvial facies association, (a) Coarse-grained channel deposits at the base of the fluvial interval. Pampa de Tril East area, (b) Coarse-grained sandstones bodies with horizontal lamination. Pampa de Tril South area, (c) Crosslaminated fine-grained sandstones intercalated between massive mudstones of the fine-grained fluvial facies association in the Chacay-Melehue sector. The coin is 3 cm in diameter, (d) Lenticular sandstone channels with large-scale inclined surfaces of the fine-grained fluvial association. Chacay Melehue area.

presence of bimodally sorted, wind-rippled sandstones. These are interpreted as aeolian sandsheets and low-relief aeolian macroforms that were formed by the localized reworking of fluvial sands during dry periods. The presence of texturally mature sand within the fluvial deposits indicates that the fluvial systems also reworked the aeolian deposits during wet periods.

Fine-grained ephemeral fluvial. This subassociation characterizes the upper portion of the Troncoso Member in the northern and NW part of the studied area. This association is characterized by a significant proportion of fine-grained deposits. Red and locally green, massive-laminated clay stones and siltstones in beds up to 3.5 m thick give this part of the succession a conspicuous red colour in outcrop. The facies show abundant desiccation cracks, rooted horizons and a mottled/blocky structure, pointing out an important degree of postdepositional modification. No identifiable trace fossils have been recognized. In addition to the mudstones, this association also comprises

very fine- to medium-grained, predominantly massive sandstones. These sandstones also contain ripple cross-lamination or small-scale trough cross-bedding in beds that range from 2 to 40 cm in thickness (Fig. 7c). Sandstone beds typically show a tabular geometry, and a sharp and horizontal lower boundary. Locally, some of the sandstones exhibit a more concave and erosive base, and a lenticular geometry. Rare, coarse-grained deposits comprise medium- to coarse-grained sandstones with medium-scale trough cross-bedding and scarce intraclast conglomerates. These coarser deposits group together in lenticular, erosive-based bodies up to 6 m thick and tens of metres wide. The bodies contain inclined large-scale stratification surfaces (sensu Bridge 1993) indicating lateral accretion (Fig. 7d). Bimodally sorted, wind-rippled very-fine- to medium-grained sandstones are also present in large-scale crossbedded sets up to 2.5 m thick or with horizontal lamination in 40 cm-thick beds. These deposits have a conspicuous sharp, horizontal base and are vertically associated with the finer grained deposits.

SEDIMENTOLOGY OF A LOWSTAND WEDGE

This sub-association is interpreted as the product of accumulation in a more distal fluvial environment relative to the coarse-grained fluvial deposits described above. The preservation of significant proportions of fine-grained material also indicates a higher accommodation rate. Deposition was mainly from unconfined floods that carried fine-grained sand and silt to the distal portions of the fluvial system. The presence of channel bodies laterally related to finegrained deposits indicates that some of these unconfined flows were probably overbank systems rather than true sheetfloods (sensu Bull 1972). Coarse-grained channel bodies may represent deposition within meandering channels that sustained flow long enough to develop point bars. Minor sandstone channels may have been more ephemeral feeder channels to the unconfined sheet deposits. Clay stones with desiccation cracks indicate the presence of shallow-water bodies that periodically dried out. Rooting and pedogenic modification indicates that the climate was, at least locally, damp enough to sustain vegetation. As with the previous sub-association, bimodally sorted sandstones are interpreted as aeolian deposits that accumulated during dryer periods. Aeolian fades association Aeolian deposits characterize the upper part of the studied interval in the southern and eastern areas. Large-scale dune features are absent in the central parts of the basin where thin windlain deposits are interbedded with distal fluvial deposits. The aeolian facies association constitutes the upper portion of the Lower Troncoso Member of the Huitrin Formation in the Pampa de Tril, Cerro Curaco, Balsa Huitrin and Cerro Rayoso areas (Fig. 5). It is also the main hydrocarbon reservoir in the subsurface and accounts for up to 35% of the total hydrocarbon production of the basin (Masarik 2002). This interval is characterized by well-sorted fine- to medium-grained sandstones with planar and trough cross-bedding on a wide variety of scales. Two separate units have been defined and these are interpreted to represent different accumulation conditions in an aeolian environment. In the Pampa de Tril area and Curaco areas (Fig. 5), the lower part of the aeolian succession is characterized by fine- to medium-grained sandstones, bimodally sorted, with horizontal lamination or small sets (<2m) of planar cross-stratification (Fig. 8a). Abundant lowangle truncation surfaces dip in a variety of directions and there is a wide spread of palaeocurrent directions.

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These deposits represent the development of very small aeolian dunes and low-relief aeolian sandsheets. Aeolian sandsheets occur in areas of low sand supply and/or where a relatively high water table and frequent flooding preclude the formation of dunes (Kocurek & Nielsen 1986; Loope & Simpson 1992). The stratigraphic position of these deposits, beneath the main aeolian succession (described below), is interpreted as representing a transitional phase from the fluvial system to the main erg. In contrast, the upper aeolian section is more widely distributed (Pampa de Tril, Cerro Curaco, Balsa Huitrin and Cerro Rayoso, Fig. 5) and characterized by the development of large-scale cross-stratified, fine- to mediumgrained sandstones in sets up to 14 m thick (Fig. 8b). These deposits show similar textural characteristics to those present in the lower section, but the intercalation of coarser grained, inversely graded laminae and fine-grained massive laminae is more common. Multiple scales of bounding surface are present between the cross-bedded sets. The predominant palaeocurrent direction is toward the NE. In those areas where the aeolian succession reaches its maximum thickness, the interval occurs as amalgamated cross-sets up to 10m thick that represent the superimposition of sinuous-crested dunes. Towards more marginal areas, this interval is dominated by superimposed sets that represent the development of complex dunes with extensive interdune areas where minor aeolian deposits were accumulated. Isolated small aeolian dunes have also been identified in marginal areas. In the southernmost parts of the study area, thin, dark mudstones and fine-grained sandstones with wavy lamination are intercalated between the large-scale cross-sets. Where the lower aeolian unit is present, the transition to the overlying large-scale dunes is across a sharp surface. Locally, this surface is overlain by minor, very-coarse-grained fluvial deposits. In those places where the lower section is missing, such as in the Balsa Huitrin area, the large-scale aeolian sets directly overlie a conspicuous truncation surface with abundant root marks (Fig. 8c). The upper boundary, at the contact with the overlying evaporites, is very irregular and shows preserved dune topography. Associated with this preserved topography is evidence for localized reworking and some soft-sediment deformation (Fig. 8d) (Stromback et al 2005). The upper aeolian unit represents the development of an erg, composed of climbing dune and draa-scale bedforms. The erg was predominantly

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Fig. 8. Aeolian facies association, (a) Small-scale cross-bedded sets with wind-ripple lamination. Pampa de tril South area. The hammer is 40 cm long, (b) Large-scale trough cross-bedded aeolian sandstones. Cerro Curaco area, person for scale, (c) Root marks at the top of aeolian cross-bedded sets. Balsa Huitrian sector. The coin is 3 cm in diameter. (d) Soft-sediment deformed sandstones at the top of the aeolian succession at the contact with the overlying evaporites. Balsa Huitrin area.

dry, although interbedded mudstones towards the margins represent wet interdune accumulation. Aeolian bedforms were predominantly complex longitudinal and transverse features with a dominant migration direction toward the NE. The sharp basal surface indicates that, although the lower aeolian unit is transitional from the fluvial deposits, there was a significant stratigraphic break between the two depositional phases that represents the development of an erosional super-surface (sensu Kocurek & Havholm 1993; Havholm & Kocurek 1994). The sub-aqueous reworking and locally preserved topography at the top of the dunes indicates that the erg was flooded fairly quickly in a similar way to that proposed for the Weissliegend deposits of the European Permian Basin (Glennie & Duller 1983; Stromback & Ho well 2002). The nature of the aeolian system and the flooding at the top of it is the subject of a companion paper in this volume (Stromback et al 2005). Evaporites The main study interval is sharply overlain by a thick succession (up to 250 m) of evaporites

belonging to the Upper Troncoso Member of the Huitrin Formation. These deposits extend beyond the study area across the entire Neuquen Basin where they overlie either the lower portion of the Huitrin Formation or, in its absence, they directly overlie the deep-marine deposits of the Upper Member of the Agrio Formation. The lower boundary of the evaporites is extremely sharp and, in the southern part of the studied area, it drapes the underlying dune topography of the main aeolian facies association (Stromback et al. 2005). In the northern area the evaporites directly overlie fine-grained deposits of the ephemeral fluvial association. The evaporitic succession is mainly composed of anhydrite and halite with thin intercalations of potassium salts (mainly sylvite). Wavy stromatolitic boundstones are also locally present at the base. These evaporites represent accumulation under hypersaline conditions in a restricted marine environment that lead to the almost complete desiccation of the basin. The thickness indicates that evaporitic conditions continued for a long period of time and the basin was continuously refreshed with marine waters (Legarreta

SEDIMENTOLOGY OF A LOWSTAND WEDGE

& Gulisano 1989; Gabriele 1992). A detailed study of these evaporites is beyond the scope of this paper. Most importantly, they represent the earliest phase of marine flooding following the major sea-level fall that led to the deposition of the coarse lowstand deposits. Palaeogeography and evolution of the lowstand wedge The study region has been subdivided into two distinct areas that are characterized by a different distribution of facies associations. Despite the differences in depositional environments the vertical evolution of these two areas is very similar. NW sector In the NW sector the succession has a variable thickness that reaches up to 150m in localities like Colorado River and Porton de Tril (Fig. 4). Although the total thickness of the studied interval is variable, the entire NW sector shows a very consistent vertical evolution. Ten to 30 m of shallow-marine carbonates sharply overlie the offshore deposits of the Upper Member of the Agrio Formation. This relationship is found in almost every locality of the NW sector and the only exception is the Porton de Tril section were the transition from the offshore marine deposits is represented by a 60 m-thick, coarseningupwards siliciclastic shallow-marine succession that is in turn overlain by a 30 m-thick succession of carbonates. The shallow-marine deposits are abruptly overlain by a coarse-grained fluvial association. This succession is characterized by up to 60 m of stacked channel units with very limited preservation of overbank deposits. The coarse-grained fluvial interval is present throughout the NW sector with the exception of the Chos Malal area, and thickens into the basin towards the north. The upper part of the studied interval is composed of deposits that belong to the finegrained ephemeral fluvial facies association. This interval also thickens towards the north, reaching a maximum of 55 m in the Colorado River area (Fig. 4). The entire sector is capped by the evaporites of the Troncoso Superior. In this area the basal contact of the evaporites is planar and sharp. This surface is interpreted as a transgressive surface. The succession in the NW sector can be summarized as a wedge that thickens towards the north and NE, with shallow-marine deposits at the base, coarse-grained fluvial deposits in the middle and a fine-grained ephemeral fluvial succession at the top, capped by evaporites.

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SE sector The succession in the SE sector differs considerably from the NW area. The studied interval is consistently thinner, with a maximum thickness of approximately 50 m occurring in the Curaco area. This compares to a maximum thickness in the NW of over 100 m (Figs 4 & 9). The SE sector also lacks shallow-marine carbonates at the base, has a considerably thinner coarsegrained fluvial succession and has a variably thick aeolian succession at the top. The depositional system in the SE sector shows a higher degree of lateral variability than the area to the north, with greater changes in both facies and facies thickness between closely spaced localities (of the order of tens of kilometres). However, as for the NW sector, it is still possible to define a common depositional history. The base of the section is marked by a sharp and locally erosive contact with the underlying offshore shales of the Agrio Formation. In a number of localities this surface is overlain by 2-20 m of shallow-marine clastic deposits. Locally, such as in the Cerro Rayoso immediately above the basal surface, the shallow-marine deposits are characterized by amalgamated HCS sandstones in a bed up to 3 m thick, with a sharp, horizontal lower boundary. In other areas such as Balsa Huitrin, the contact is dominated by the presence of a layer with gutter casts (Fig. 6a). The shallow-marine deposits of the SE sector are predominantly clastic, off shore-transition-lower-shoref ace deposits. No uppershoreface deposits were observed. In contrast to the NW sector, bioclastic sandstones are common and shallow-marine carbonate deposits are absent. Above the shallow-marine deposits, or directly on top of the offshore marine deposits (e.g. at Cerro Curaco or Pampa de Tril), is a sharpbased succession of coarse-grained fluvial deposits. These fluvial deposits typically sit within large-scale scours, cut into both the shallowand deep-marine deposits beneath them. These scours account for much of the marked thickness changes (0-20 m) in this unit. The coarsest grained fluvial deposits typically sit at the base of the successions that are comprised of amalgamated channel deposits with very rare interbedded overbank deposits. Upwards the succession becomes more sandstone dominated, channels become less confined and are interbedded with aeolian sandsheet deposits. The upper portion of the studied interval in the SE sector is characterized by the development of a thick aeolian facies association that is completely absent in the northern area. This part of the succession may reach up to 35 m

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Fig. 9. General distribution and correlation of the facies associations across the entire study area.

thick and has a slightly more widespread distribution than the underlying portion. The aeolian deposits are present in all of the studied sections. They overlie the fluvial facies association in the Cerro Rayoso, Cerro Curaco and Pampa de Tril areas, and directly overlie the shallow-marine facies association in the Balsa Huitrin and Pampa de Tril (E) areas. Differences in the aeolian bedforms were used to define two different aeolian systems separated by a major bounding surface. The lower aeolian unit is characterized by small bedforms and abundant sandsheet deposits, whilst the upper is comprised of large-scale dunes. Overall, the transition from the shallow-marine and/or fluvial deposits to the upper aeolian unit represents a drying-upwards sequence with an increase in sand supply (or sand availability) (Kocurek & Lancaster 1999). As in the NW sector, this interval is overlain by the evaporites of the Troncoso Superior Member. In contrast to the sharp planar surface seen above the distal fluvial deposits in the NW, in the SE the contact is more complex. The geometry of the surface is highly irregular, preserving in some degree the aeolian topography that was present prior to the flooding (Stromback el al 2005). Where the aeolian deposits are absent and the evaporites directly overlie shallow-marine deposits this upper boundary is characterized by a transgressive lag of coarse-grained bioclastic sandstone.

Sequence stratigraphy While it is not physically possible to trace surfaces between the various outcrop sections, a correlation based on changes in relative sea-

level has been proposed. Broadly, the study interval represents a lowstand wedge, overlying the highstand deposits of the Agrio Formation and capped by a transgressive systems tract represented by the evaporites of the overlying Troncoso Superior. The following section describes the internal geometry of the lowstand wedge with respect to a number of internal key surfaces, controlled by changes in relative sea level. These surfaces include (Figs 10 & 11): the regressive surface of marine erosion; the main sequence boundary; the transgressive surface at the top of the wedge. The regressive surface of marine erosion — falling-stage systems tract The deposits of the upper Agrio Formation represent the HST of a third-order sequence deposited within a ramp-type basin (Spalletti et al. 200 Ib). This sequence had a far greater extent than that of the main study interval. Towards the basin margins the upper portion of the Agrio Formation is characterized by a progradational parasequence set (Spalletti et al. 200la). Whilst a certain degree of gradual shallowing is apparent in the upper parts of the Agrio Formation within the study area, the abrupt transition from offshore facies to shallow-marine shoreface deposits records a relative sea-level fall. In the SE sector the shallow-marine interval overlying this transition is less than 10m thick. The basal surface is marked by the deposition of 3 m of amalgamated HCS beds in the Cerro Rayoso area, by a 7 m shoreface with well-developed E-W-trending gutter casts in the Balsa

Fig. 10. Chronostratigraphic chart and interpreted sequence stratigraphic scheme for the studied interval.

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Fig. 11. Interpreted relative sea-level/base-level evolution for the studied interval.

Huitrin area and by an abrupt change to offshore transition facies in the Pampa Tril sections. All of these successions suggest a rapid lowering of the storm wave base (Fitzsimmons & Johnson 2000) as a response to relative sea-level fall. These deposits can consequently be regarded as a forced regressive systems tract (FRST, Hunt & Tucker 1995) or a falling-stage systems tract (FSST, Flint & Nummedal 2000) and are bounded at the base by a regressive surface of marine erosion and at the top by the subaerial

unconformity (sequence boundary) (Figs 10 & 11). In the NW sector, the same stratigraphic interval includes a thick (60 m), but laterally very restricted, siliciclastic shallow-marine succession that is only present in one area (Porton de Tril) (Fig. 9). Carbonate deposits are present throughout the NW sector. At Porton de Tril the carbonates overlie the thick shoreface succession, indicating that the siliclastics predated the carbonates. The greater thickness of this most

SEDIMENTOLOGY OF A LOWSTAND WEDGE

northerly siliciclastic succession and the presence of a complete progradational succession (up to upper shoreface) represent different accumulation conditions in the north. The northern sectors appear to have been dominated by higher sediment-supply conditions also manifested by the abundance of soft-sediment deformation structures. With the exception of the Porton de Tril area the carbonates in the northern sector lie directly on the offshore deposits of the Agrio Formation. These carbonate deposits imply a complete change in accumulation conditions from an open-marine setting with considerable siliciclastic supply to a carbonate ramp. The predominance of algal boundstones in the central part of the basin and the absence of other marine fossils may indicate an increase in salinity related to isolation from the Pacific Ocean (Gutierrez Pleimling 1991). This change is interpreted to be a result of relative sea-level fall coupled with a complex basin topography that isolated parts of the basin from clastic supply (Figs 10 & 11). Master sequence boundary and the overlying lowstand systems tract The accumulation of the shallow-marine facies described in the previous section does not represent the maximum stage of marine regression. Throughout much of the studied area the shallow-marine deposits are incised by a major subaerial erosion surface that is overlain by coarse-grained fluvial deposits (Fig. 9). In other parts of the study area the same surface is characterized by a sharp transition from lowershoreface to aeolian deposits, also recording an abrupt lowering of sea level. This surface represents the lowest point on the relative sea-level curve and is regarded as a sequence boundary (sensu Hunt & Tucker 1992, 1995). This surface is the most significant of a number of surfaces that represent relative drops in sea level; it is consequently termed the master sequence boundary (MSB, Figs 10 & 11). The MSB is followed by the development of non-marine deposits in the whole study area, after accommodation is created once again in the basin. The lateral transition to non-marine evaporates in the centre of the basin, north of the main study area (Veiga & Rossi 1992), suggests that complete desiccation of the basin occurred. This desiccation was associated with disconnection of the basin from the palaeoPacific Ocean, a situation proposed for other lowstand periods in the Mesozoic history of the Neuquen Basin (Mutti et al 1994; Legarreta 2002).

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The MSB is overlain by non-marine deposits that typically pass from coarse-grained fluvial facies upwards into aeolian deposits in the SE and fine-grained fluvial deposits in the north. Although in individual sections this transition may look sharp, local aeolian reworking of the coarse fluvial deposits and the presence of small aeolian bedsets at the base of the aeolian section indicate that this change was gradual and does not represent a single, basinwide event, but rather a gradual decline in fluvial activity and an increase in aeolian deposition. This may be the result of a gradual climate change towards dryer conditions with an increase in wind strength. However, this vertical succession may also be the result of a gradual watertable lowering that progressively increased the amount of sand available to be transported and accumulated by winds that were already active when the system was dominated by fluvial processes, as shown by local aeolian reworking of fluvial material. The lateral correlation of the fluvial deposits with non-marine evaporites indicates that the basin was already desiccated; therefore this water-table drop may have just been the result of the climatic effect of evaporation of remaining water in a land-locked basin and may have no relationship to sea-level fluctuations. Previous workers (Legarreta 2002) have proposed that this period of aeolian activity represents the initial stage of the transgressive systems tract. During transgressive periods, and even under limited sand supply, relative sealevel rise will generate a rise in water table allowing the preservation of previously accumulated aeolian deposits (cf. Kocurek et al. 2001). The presence of isolated, non-climbing aeolian bedforms (Fig. 12) indicates limited accommodation that could be promoted by a low water table during this lowstand period. The subsequent flooding and accumulation of marine evaporites on top of this aeolian succession represents an 'instantaneous' rise in water table (sensu Kocurek 1999) that promoted the preservation of dune topography and isolated dunes. All this suggests that these aeolian deposits were accumulated during a period of relatively low water table (lowstand) rather than during a transgressive event.

Trangressive surface - transgressive systems tract The top of the studied interval is characterized by the accumulation of a thick package of marine evaporites that are present across the entire basin (the Upper Troncoso Member of the

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Fig. 12. Isolated aeolian bedforms in the Pampa de Tril East area. Note the preserved dune topography at the contact with the overlying evaporites.

Huitrin Formation) (Fig. 9). The accumulation of marine evaporites on top of a non-marine succession represents a relative rise in sea level and the flooding of at least part of the basin. The transgression is associated with local marine reworking of aeolian deposits and the formation of a transgressive surface of erosion (TSE, Figs 10 & 11), although aeolian topography of up to 30 m is preserved in many areas. This topography and the presence of preserved, non-climbing bedforms in the previous erg margin localities indicates a rapid transgression (Stromback et al 2005). In some areas the complete non-marine succession is absent and the transgressive surface lies directly on top of shallow-marine deposits of the falling-stage systems tract and is characterized by a lag of fragmented bioclasts. Marine reworking is least in the NW sector indicating that this area was flooded most rapidly or under lower wave energy conditions. The flooding of the basin is interpreted as a rapid event, potentially associated with sudden re-opening of the basin to the palace-Pacific. Controls on the formation of the lowstand wedge: discussion The deposits described above have been interpreted as the result of a second-order fall and subsequent rise in relative sea level (Legarreta & Gulisano 1989). The Mesozoic fill of the Neuquen Basin contains the deposits of several high-magnitude, low-frequency cycles that

resulted, at lowstand, in the deposition of nonmarine deposits in the centre of the predominantly marine basin (Legarreta 2002). Although the study interval is broadly similar to these other cycles, there are also some significant differences that warrant discussion. These differences are characterized by the overall dryingupwards trend within the wedge and the nature of the sediments above the lowstand deposits. The vertical trend in facies within the lowstand deposits shows an upwards increase in the relative importance of aeolian activity. The lower deposits are fluvial dominated with only limited aeolian reworking. The lower fluvial-dominated succession passes up into a mixed succession of sandy fluvial deposits interbedded with small aeolian dune deposits, which in turn are overlain by sediments of the main erg. Overall, the succession represents a decrease in fluvial activity and an increase in aeolian processes. Drying-upwards cycles are common within semi-arid continental successions (Clemmesen et al. 1989; George & Berry 1993; Howell & Mountney 1997) and are typically interpreted to result from allocyclic changes in climate. These changes may have been caused either by long-term global climate changes or by more localized changes within the basin. Sarnthein (1978) demonstrated a link between ice-sheet development and the compression of the tradewind belts that lead to increased aeolian activity within mid-latitude desert systems. Such a mechanism was proposed by Howell (1992) as a control on drying-upwards cycles within the Permian Rotliegend deposits of the North Sea

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Fig. 13. Evolution of the SW margin of Gondwana during the Cretaceous and implication for basin configuration, (a) Normal situation before the accumulation of the studied interval. Note that both highstand and lowstand sea levels allow a complete connection between the basin and the Pacific Ocean, (b) Situation during the accumulation of the studied interval. Note that the combination of the arc growth and a relative sea-level fall might have disconnected the basin completely from the open-marine setting during the stage of maximum retreat, (c) Situation during the accumulation of the overlying evaporites. The growth in the arc might have prevented the complete connection with the open ocean leading to resticted marine back-arc basin. Modified from Ramos (1999).

(see also George & Berry 1993; Ho well & Mountney 1997). Uplift in the Andes, creating a rainshadow effect, may equally have caused a drying in the climate. Alternatively, the drying

upwards may simply result from autocyclic developments within the erg, which, given a steady supply of sediment, may have caused the dune system to climb above the water table.

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Also, the desiccation of the basin due to the evaporation of remaining water after the complete disconnection from the Pacific Ocean may have resulted in an overall water-table lowering leading to an increase in sand availability for aeolian accumulation within the lowstand succession. Two previous second-order sea-level falls occurred in the Valanginian and Hauterivian of the Neuquen Basin (Veiga et al. 2002; Schwarz & Howell 2005). After both of these sea-level falls there was a return to openmarine conditions. In the study interval, however, the lowstand wedge is overlain by a thick Aptian-Albian succession of evaporites. This implies that the reconnection to the Pacific Ocean was only partial and the basin became a restricted marine setting. The lack of openmarine circulation may have resulted either from lower global sea levels or from a change in the nature of the volcanic arc. As eustatic sea levels were relatively high during the Lower Cretaceous (Haq et al. 1988) it is concluded that the isolation of the basin resulted from the vertical growth of the volcanic arc. This growth corresponds to a change in the dynamics of the continental margin that has been proposed for the end of the Early Cretaceous in the active margin of western South America (Ramos 1988, 1999; Cobbold & Rosello 2003). This change represents the transition from an extensional retro-arc environment (where the rate of subduction exceeded the rate of plate convergence) to an Andean-type margin (where convergence rate was faster than subduction rate); with a transitional stage during the Aptian where subduction and convergence rates were similar. During this transitional phase, uplift in the western margin of the basin may have uplifted the threshold between the Neuquen Basin and the Pacific Ocean, restricting its connection (Fig. 13). It is unclear whether this growth caused the initial desiccation of the basin or simply prevented the redevelopment of openmarine conditions after the relative sea-level fall. However, this additional evidence for uplift within the volcanic arc indicates a basin-wide, rather than global, control on the dryingupwards cycle observed within the succession. Conclusions The Lower Cretaceous deposits of the Neuquen Basin include evidence for large-scale relative sea-level changes. The changes resulted in the formation of shallow clastic and carbonate environments on top of the offshore marine Agrio Formation. Further sea-level fall led to the deposition of an arid, fluvial-aeolian

drying-upwards succession within the study area. This succession is capped by evaporites that represent the early stages of the subsequent transgression. Fully marine conditions are not re-established in the basin. This large-scale change in the depositional setting is attributed to uplift of the Andes and the basin's transition from a back-arc to a foreland-type system. This research has been funded by the National Research Council of Argentina (CONICET) through a Postdoctoral Research Fellowship and by the Ministry of Education, Science and Technology of Argentina, through the IM40 Grant #34. G. Plint and C.O. Limarino are thanked for their constructive reviews of the original manuscript.

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HUNT, D. & TUCKER, M.E. 1995. Stranded parasequences and the forced regressive wedge systems tract: deposition during sea level fall - reply. Sedimentary Geology, 95, 147-160. KOCUREK, G. 1999. The Aeolian Rock Record. In: GOUDIE, A.S., LIVINGSTONE, I. & STOKES, S. (eds) Aeolian Environments, Sediments and Landforms, John Wiley & Sons, Ltd., 239-259. KOCUREK, G. & HAVHOLM, K.G. 1993. Eolian sequence stratigraphy - a conceptual framework. In: WEIMER, P. & POSAMENTIER, H. (eds) Recent Advances and Applications of Siliciclastic Sequence Stratigraphy. AAPG Memoirs, 58, 393-409. KOCUREK, G. & LANCASTER, N. 1999. Aeolian syste sediment state: theory and Mojave Desert Kelso dune field example. Sedimentology, 46, 505-515. KOCUREK, G. & NIELSON, J. 1986. Conditions favour able for the formation of warm-climate aeolian sand sheets. Sedimentology, 33, 795-816. KOCUREK, G., ROBINSON, N.I. & SHARP, J.M., JR. 2001. The response of the water table in coastal aeolian systems to changes in sea level. Sedimentary Geology, 139, 1-13. LEGARRETA, L. 1985. Andlisis Estratigrdfico de la Formacion Huitrin (Cretacico Inferior), Provincia de Mendoza. PhD Thesis, Universidad de Buenos Aires. LEGARRETA, L. 1986. Litogenesis de las secuencias depositacionales de la Formacion Huitrin (Cretacico Inferior), Provincia de Mendoza. In: SPALLETTI, L.A. (ed.) Primera Reunion Argentina de Sedimentologia. Resumenes Expandidos, La Plata, 173-176. LEGARRETA, L. 2002. Eventos de desecacion en la Cuenca Neuquina: depositos continentales y distribution de hidrocarburos. In: V° Congreso de Exploracion y Desarrollo de Hidrocarburos. Tremp, Spain. Technical works on CD. LEGARRETA, L. & GULISANO, C.A. 1989. Analisis estratigrafico secuencial de la Cuenca Neuquina (Triasico Superior-Terciario inferior). In: CHEBLI, G. & SPALLETTI, L.A. (eds) Cuencas Sedimentarias Argentinas. Universidad Nacional de Tucuman, Serie Correlation Geologica, 6, 221-243. LEGARRETA, L. & ULIANA, M. 1999. El Jurasico y Cretacico de la Cordillera Principal y la Cuenca Neuquina. 1. Facies Sedimentarias. In: CAMINOS, R. (ed.) Geologia Argentina. Instituto de Geologia y Recursos Minerales, Anales, 29, 399-416. LEGARRETA, L., ULIANA, M.A., LAROTONDA, C.A. MECONI, G.R. 1993. Approaches to nonmarine sequence stratigraphy - Theoretical models and examples from Argentine basins. In: ESCHARD, R. & DOLIGEZ, B. (eds) Subsurface Reservoir Characterization From Outcrop Observations. Editions Technip, Paris, 125-143. LOOPE, D.B. & SIMPSON, E.L. 1992. Significance of thin sets of eolian cross-strata. Journal of Sedimentary Petrology, 62, 849-859. MASARIK, M.C. 2002. Los Reservorios de las Formaciones Agrio y Huitrin: Introduction. In: SCHIUMA, M., HlNTERWIMMER, G. & VERGANI,

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SPALLETTI, L.A., POIRE, D.G., PIRRIE, D., MATHEOS, S. & DOYLE, P. 200la. Respuesta sedimentologica a cambios en el nivel de base en una secuencia mixta clastica-carbonatica del Cretacico de la Cuenca Neuquina, Argentina. Revista de la Sociedad Geologica de Espana, 14, 57—74. SPALLETTI, L.A., POIRE, D.G., SCHWARZ, E. & VEIGA, G.D. 200Ib. Sedimentologic and sequence stratigraphic model of a Neocomian marine carbonate-siliciclastic ramp: Neuquen Basin, Argentina. Journal of South American Earth Sciences, 14, 609-620. STROMBACK, A. & HOWELL, J. 2002. Distribution and reservoir properties of the soft sediment deformed Weissliegend the UK, Southern North Sea. Petroleum Geoscience, 8, 237-249 STROMBACK, A., HOWELL, J.A. & VEIGA, G.D. 2005. The transgression of an erg - sedimentation and reworking/soft-sediment deformation of aeolian facies: the Cretaceous Troncoso Member, Neuquen Basin, Argentina. In: VEIGA, G.D., SPALLETTI, L.A., HOWELL, J.A. & SCHWARZ, E. (eds) The Neuquen Basin: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 163-183. VEIGA, R. & Rossi, G. 1992. Analisis sedimentologico del Miembro Troncoso Inferior (Formacion Huitrin) en el ambito de la Sierra de Reyes. Dpto. Malargiie, provincia de Mendoza. In: IV Reunion Argentina de Sedimentologia, Vol. I. La Plata, Argentina, 71-78. VEIGA, R. & VERGANI, G.D. 1993. Depositos de nivel bajo: nuevo enfoque sedimentologico y estratigrafico del Miembro Avile en el Norte del Neuquen. Argentina. XII0 Congreso Geologico Argentino y 11° Congreso de Exploracion de Hidrocarburos, Actas, I, 55-65. VEIGA, G.D., SPALLETTI, L.A. & FLINT, S. 2002. Aeolian/fluvial interactions and high resolution sequence stratigraphy of a non-marine lowstand wedge: The Avile Member of the Argio Formation (Lower Cretaceous) in central Neuquen Basin, Argentina. Sedimentology, 49, 1001-1019. VERGANI, G.D., BARRIONUEVO, M., SOSA, H. & PEDRAZZINI, M. 2001. Analisis estratigrafico secuencial de alta resolution en las Formaciones Agrio y Huitrin en el Yacimiento Puesto Hernandez, Cuenca Neuquina, Argentina. Boletin de Informaciones Petroleras, 67, 76-87. VERGANI, G.D., SELVA, G. & BOGGETTI, D.A. 2002. Estratigrafia y modelo de facies del Miembro Troncoso Inferior, Formacion Huitrin (Aptiano), en el noroeste de la Cuenca Neuquina, Argentina. XV° Congreso Geologico Argentino (El CCalafate), Actas, I, 613-618. VERGANI, G.D., TANKARD, A.J., BELOTTI, HJ. & WELSINK, HJ. 1995. Tectonic evolution and paleogeography of the Neuquen Basin, Argentina. In: TANKARD, A.J., SUAREZ SORUCO, R. &. WELSINK, HJ (eds) Petroleum Basins of South America. AAPG Memoirs, 62, 383-402.

The transgression of an erg - sedimentation and reworking/ soft-sediment deformation of aeolian fades: the Cretaceous Troncoso Member, Neuquen Basin, Argentina ANNA STROMBACK1'2, JOHN A. HOWELL1'3 & GONZALO D. VEIGA1'4 1

Stratigraphy Group, Department of Earth Sciences, University of Liverpool, Brownlow Street, Liverpool L69 3GP, UK ^Present address: Statoil ASA, Forus Vest, 4035 Stavanger, Norway (e-mail: [email protected])

3

Present address: Centre for Integrated Petroleum Research, University of Bergen, Allegaten 41, Bergen N-5007, Norway

4

Centro de Investigaciones Geologicas, Calle 1 #644. B1900TAC La Plata, Argentina Abstract: The Cretaceous Troncoso Inferior Member of the Huitrin Formation comprises fluvial and aeolian facies that form a drying-upwards succession within the Neuquen Basin. The basal fluvial sandstones were deposited as braided river deposits and lie unconformably on top of either the deep-marine Agrio Formation or, locally, the shallow-marine Chorreado Member (Huitrin Formation.). In places, the fluvial sandstones are interbedded with remnants of aeolian deposits recording an arid environment and ephemeral flows. In the study area the upper section is predominantly aeolian and was controlled by northerly winds with both linear and transverse dune types being deposited. The depositional system was rapidly flooded and dune topography (relief ranging between 2 and 35 m) was preserved on its top surface. In addition to dune topography, the Troncoso dunes also show evidence of reworking and in situ soft-sediment deformation related to the flooding. The principal aim of this paper is to document the soft-sediment deformation and preservation of topography associated with the flooding of the dune field. Within the soft sediment deformed and reworked sediments at the top of the Troncoso Inferior Member spatial and temporal relationships indicate that they formed in a specific sequence. Initially, water-escape processes created convolutedly folded and dish structures that were concentrated in areas of slightly higher preserved dune topography. Secondly, the convolutedly folded and dish facies were eroded and reworked by wave undercutting and migrating three-dimensional dunes in a shallow-marine environment. This subaqueous reworking resulted in an interbedded massive and cross-stratified unit. With further deepening of the water, the topography became stabilized and the uppermost part of the interval (0.1-0.3 m) was reworked by waves across most of the basin. In the topographic lows between dunes, liquefaction-induced sediment gravity flows deposited massive-flatlaminated facies. The reworked and soft-sediment deformed aeolian dune topography is overlain by the evaporites of the Troncoso Superior Member. The distribution of flood-related facies and the amount of preserved dune topography (2-35 m) indicates that the transgression must have been rapid but of low energy.

The aim of this study is to address geometric and process aspects of reworked and soft-sediment deformed facies in a drowned aeolian system. The Late Aptian Troncoso Inferior Member of the Neuquen Basin (Argentina) includes aeolian deposits towards its top, which are overlain by carbonates and evaporites. The excellent outcrops east of Chos Malal (Fig. 1) provide an

ideal opportunity to study flooded dune systems. Depositional super-surf aces (sensu Kocurek & Havholm 1993) created by the flooding of dune systems have been recognized previously in both outcrop and subsurface studies (e.g. Vincelette & Chittum 1981; Eschner & Kocurek 1986, 1988; Mountney^ al. 1999). Preserved dune topography is also seen in other

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 163-183. 0305-8719/05/$15.00 © The Geological Society of London 2005.

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Fig. 1. Location of the study area within the Neuquen Basin. Modified after Uliana & Legarreta (1993). The numbers marked on the location maps to the right indicate the logs in Figure 3.

systems such as the Permian RotliegendWeissliegend system (UK, southern North Sea) (Glennie & Buller 1983; Ho well & Mountney 1997; Stromback & Ho well 2002), the Jurassic Entrada Sandstone (SW USA) (Benan & Kocurek 2000) and the Page Sandstone (SW USA) (Blakey et al 1996). Within the Troncoso the architecture of the aeolian and fluvial sequence was studied in order to understand the depositional system that were active immediately prior to the transgression. This is summarized within this paper and addressed more fully in an accompanying paper (Veiga et al. 2005). The main focus of this study is the preserved dune topography beneath the flooding surface, and the distribution of reworked and soft-sediment deformed facies. The study area is situated in the NW part of the Neuquen Basin in Argentina (Fig. 1). This basin contains large amounts of Argentina's hydrocarbon reserves, and the Troncoso is an important reservoir interval (Fig. 1) (Uliana & Legarreta 1993). Uliana et al. (1975), Legarreta (1985) and Legarreta & Uliana (1991) have previously studied the Huitrin Formation. Veiga et al. (2005) provide a detailed study of the entire Troncoso Inferior Member; however, this is the first detailed study of the upper soft-sediment deformed and reworked part of the system.

Basin history and regional stratigraphy The Neuquen Basin originated during the Late Triassic due to the extensional collapse of an

Upper Palaeozoic thickened crust (Franzese & Spalletti 2001). It started as a series of NWSE-trending rifts in an intra-arc setting on the South American foreland. During the Early Jurassic, regional thermal subsidence (Legarreta & Gulisano 1989; Legarreta & Uliana 1991) related to post-extensional cooling of the extended lithosphere (Uliana & Legarreta 1993) resulted in the development of a wide marine embayment on the SW margin of Gondwana. During the Jurassic and Cretaceous the Neuquen Basin was connected to the Pacific by a narrow seaway through openings in the magmatic arc to the west (Uliana & Legarreta 1993). The sedimentation in the basin was controlled both by eustatic sea-level changes in the main Pacific and by tectonic uplift of the arc. The basin was periodically isolated from the open ocean to the west by eustatic sea-level falls. The sediments that filled the basin were sourced predominantly from the SE and prograded towards the NW (Legarreta & Uliana 1991). During the Early Aptian a major fall in sea level resulted in the deposition of the continental deposits of the Huitrin Formation on top of the marine deposits of the Agrio Formation (Veiga et al 2005). Towards the centre of the basin, the Huitrin Formation includes ephemeral fluvial deposits, a saline mud flat assemblage and evaporites (Legarreta 1985). Towards the basin margins fluvial and aeolian sediments were deposited. Flooding in the Late Aptian submerged the aeolian/fluvial facies and a closed hydrological regime resulted in the deposition

THE TRANSGRESSION OF A CRETACEOUS ERG of anhydrite. Towards the margins of the basin the anhydrites are nodular with an algal matrix, whilst towards the basin centre there is 'varved' anhydrite and bedded halite (Legarreta & Uliana 1991). In the study area, the base of the Troncoso Member overlies marine deposits of the Chorreado Member (Huitrin Formation) and the Agrio Formation (Fig. 2). The Troncoso Member is divided into two intervals: the Troncoso Inferior Member, which is composed primarily of siliciclastic material, and the Troncoso Superior Member, which is dominated by evaporites (Veiga & Rossi 1992) (Fig. 2). This study concentrates on the Troncoso Inferior Member that generally has a lower fluvial and an upper aeolian portion. Small-scale dune deposits dominate the lower part of the aeolian succession, while the upper part is characterized by larger scale aeolian bedforms (Fig. 3). The

Fig. 2. The Cretaceous stratigraphy of the northern Neuquen Basin with the studied interval highlighted. Modified after Gulisano & Gutierrez Pleimling (1995).

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thickness of both the fluvial and aeolian deposits vary considerably: north and west of the study area no aeolian facies have been observed, whilst 50 km east of the study area subsurface data record a thickness of up to 100 m of Troncoso Inferior Member deposits (Comeron 1990). The upper surface of the Troncoso Inferior Member is irregular and represents preserved dune topography that varies in height from 2 to 35 m (Fig. 3). At the top of the aeolian interval reworked and/or soft-sediment deformed facies of varying types and thickness are present. Within the study area, the evaporites of the Troncoso Superior Member are mainly white, laminated anhydrites, in a 5-10 m-thick unit.

Facies associations The Troncoso Inferior Member within the study area is composed of the fluvial (FA-F), aeolian (FA-A), soft-sediment deformed (FA-SSD) and reworked (FA-R) facies associations. This study concentrates on the aeolian deposits that were present prior to the flooding, and the reworked and soft-sediment deformed facies that formed when the basin was flooded. A more in-depth discussion of the Troncoso facies and their broader distribution can be found in Veiga et al. (2005). Fluvial facies association (FA-F) The Troncoso Member includes a variety of different fluvial deposits. These are considered in some detail by Veiga et al. (2005). The following brief description refers only to the fluvial deposits that occur within the study area. They are discussed because they form an integral part of the overall drying-upwards succession that comprises the member. Within the study area the base of the Troncoso Member is an erosive unconformity that is overlain by conglomeratic and sand-rich, trough cross-bedded deposits in large cross-bed sets up to 2 m thick (Fig. 4a). Clasts include mudstones from the underlying Agrio Formation and limestones. Overall, there is an upwards decrease in cross-bed thickness and mean grain size. The upper part of the fluvial succession is comprised largely of reworked aeolian sand, and pebble clasts are rare-absent. The units are interbedded with minor silty intervals, small aeolian dune and aeolian sandsheet deposits. These fluvial deposits provide evidence of an ephemeral, braided channel belt. The massive and cross-bedded sets with conglomeratic base were formed by erosive bedload processes within fluvial channels (Miall 1988). On the floodplains adjacent to the channels, crevasse

Fig. 3. Logs correlated across the study area. Log numbers 1-5 are from Pampa de Tril, and 6 and 7 from Curaco. SB, sequence boundary; FS, flooding surface.

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Fig. 4. Troncoso Member facies. (A) Within the study area the base of the Troncoso is represented by massive and cross-bedded fluvial facies lying directly on the offshore marine deposits of the Agrio Formation. (B) & (C) Large-scale dune facies. (D) Small-scale dune facies. (E) Aeolian sandsheet association with horizontal wind-ripple laminated facies. Scale: the lens cap is 5 cm, the hammer is 30 cm and the pole 120 cm (each segment is 10 cm).

splays deposited plane-laminated and rippled sandstones during ephemeral sheet flood events (Williams 1971). The well-sorted nature of the fluvial deposits in the upper part of the fluvial succession may suggest that parts of the fluvial system were active during aeolian sedimentation and locally reworked the texturally mature dune deposits. Aeolian facies association (FA-A) Aeolian deposits lie directly above the fluvial facies in the Troncoso succession (Fig. 4b-e).

The basal aeolian sandstones are fine- to coarse-grained and generally poorly sorted. They are finely laminated (Fig. 4e), although thicker lenses of coarse-grained material are seen locally. Individual units are generally not more than 0.5 m thick and are often of limited lateral extent (Fig. 4d). Higher up in the aeolian succession, small-scale trough cross-bedded sets (< 1 m) also of limited lateral extent (metres to tens of metres) occur. These are composed of fine- to coarse-grained sandstones that are moderately well rounded, and moderate to well sorted. The foresets dip at 10°-30°

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towards the NE (in Pampa de Tril) and ENE (in Curaco). They are coarser grained towards the base of the sets where the coarse laminae pinch out and granule lenses are commonly observed. Locally (in Pampa de Tril), the small-scale, cross-bedded sandstones are intercalated with fluvial deposits. Higher in the section, the cross-beds become larger (up to 7 m) and contain fine- to mediumgrained sandstones. The sandstones of these large cross-sets are pink-grey in colour, and are generally very well sorted and contain wellrounded grains (Fig. 4b). Individual crosslaminae are a few centimetres thick and are separated by fine-grained laminae. The crossbedding dips at 15°-35° (becoming shallower at the base of sets) and are directed towards the NNE in Curaco, but have a more widespread (northerly) to bimodal directions (towards the north and east) in the Pampa de Tril. Dune cross-sets truncate each other (Fig. 4c), and the base of the sets are locally poorly sorted and coarser grained, and form tabular and finely laminated intervals that can be up to 0.3 m thick. These intervals locally contain lenses of coarse-grained sand and granules, and have bimodal lamination. This texture is similar to the tabular sets that are seen at the base of the aeolian succession but here the deposits are both thinner and laterally less extensive. The basal portion of the aeolian deposits was mainly formed by wind-ripple migration on a sandsheet or sand flat. Subcritically climbing wind ripples formed the bimodal lamination at the basal part of the succession. The coarse material within the facies probably originates from the reworking of the underlying fluvial material. Aeolian sand sheets formed because there was either insufficient loose fine-grained sand or because the wind regime was unsuitable to develop dunes. Kocurek & Nielson (1986) further suggest that a high water table plays a major role in the formation of sandsheets as the water table binds the sediment to the surface due to capillary water tension, reducing sand availability (see Kocurek & Lancaster 1999). Other factors that might also favour sandsheet formation are surface cementation, periodic flooding, a significantly coarser grain size and the presence of vegetation (Kocurek & Nielson 1986). The small- and large-scale cross-bedded sets formed by grain-flow/fall processes on the lee-side of dunes (Hunter 1977). Grain-flow processes formed the coarser grained laminae, and grain-fall processes the finer grained laminae. Subcritically climbing ripples deposited windripple laminated and, occasionally, bimodally

sorted sandstones at the base of sets and overlying second-order bounding surfaces (called superimposition surfaces: Kocurek 1981) in the lower part of the succession. The coarser grained character of the cross-beds at the base of the succession indicates that it was deposited in an environment with insufficient fine-grained sand to build large dunes. Finer grained sand may have been trapped in adjacent areas, for example in wet/damp interdune areas, or may have been winnowed by strong winds and carried elsewhere. In addition, there may have been a lack of finer grained sand when the dunes were being built. The granular material that is locally observed at the base of aeolian foresets indicates a time of deflation and winnowing of material deposited during flash flood events. Higher up in the succession, the sets become thicker and the predominance of wind-rippled strata on bounding surfaces with lack of evidence for a water-table influence (adhesion structures etc.) indicates a dry aeolian system (Kocurek & Havholm 1993). Superimposition surfaces are absent and, instead, subsets are separated by widely spaced third-order bounding surfaces that may indicate changes in wind direction (Kocurek 1996).

Soft-sediment deformed facies association (FA-SSD) This facies association comprises mainly palegrey sandstone with the same textural properties as the large-scale aeolian cross-bedded deposits. Grain-fall and grain-flow laminae of aeolian origin are folded to varying degrees of complexity into structures such as convolute laminae, wavy subparallel bedding, cone-shaped diapirs and broad synclines (Fig. 5c, d). Bed boundaries within the units are diffuse. The thickness of units with convolute folding varies from single cross-beds (cm) to whole cross-sets up to several metres in thickness. Some examples of folds are slightly overturned, although no predominant sense of shear was observed. Within units, the amount of deformation decreases downwards and the folded strata pass gradually into the underlying aeolian strata. Ball-andpillar-like structures with concave-upwards, subparallel bedding are locally observed. Interbedded and in association with the convolutedly folded beds are dish structures. The individual dishes are 0.1-0.3 m wide and have a concaveup form (Fig. 5e). The dark rusty coloured dishes are finer grained than the rest of the generally massive sandstone. The contact between the soft-sediment deformed material and the

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Fig. 5. Reworked and soft-sediment deformed facies association. (A) Massive-flat-laminated facies (section c. 1 m high) contain sets with (B) rippled top surfaces, (C) convolute folded facies in the form of a cone-shaped diapir; and (D) more complex folded pattern, (E) dish facies, with gently curved dishes about 15 cm wide, (F) interbedded massive and crossbedded facies with erosive contact to underlying aeolian large-scale dune facies, (G) well-cemented wave-rippled facies forms the top of the succession and is locally comprised of (H) cross-stratification and lies directly below the Troncoso Superior Member. Scale: the lens cap is 5 cm, the hammer is 30 cm and the pole 120 cm (each segment is 10 cm).

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underlying aeolian deposits is gradational, and a thin interval (c. 0.5 m) of massive sandstone is often observed in between the fades associations. The aeolian deposits that underlie the soft-sediment deformed sandstones are locally slightly modified and show structures such as wavy diffuse subparallel laminae and simple overturned laminae. The soft-sediment deformed material is the result of rapid upwards escape of water and/or air associated with pressure changes within the dunes resulting from flooding (Lowe 1975). The fluid and gas movements folded and deformed existing laminations. As the dunes became rapidly saturated, trapped pockets of air escaped upwards. Loading of water-saturated sand associated with further flooding resulted in the upwards migration of water and the saturated laminations deformed plastically (Lowe 1975). Further deformation results in loss of most of the original fabric and the formation of complexly folded strata. Localized massive deposits result from the total loss of cohesive strength within the sediment caused either by rapid movements of fluid or seepage of water through the sediment (Lowe 1975). The dish structures form when grain-supported and more compacted sand sinks into less dense fluidized zones. The sinking sand replaces material carried up by fluidization. The upwards-migrating flow carries elutriated fine-grained material that attaches to the base of the sinking bodies which form the side of the flow path for the escaping fluid-sediment mixture (Lowe & LoPiccolo 1974; Rautman & Dott 1977). This process is most common when the fluidization or liquefaction of the sediment is unevenly distributed. Reworked fades association (FA-R) This facies association has a range of depositional styles. In the Pampa de Tril area reworked material is deposited as massive-flat-laminated sandstones (up to 15 m thick) that have the same textural properties as the underlying aeolian dune cross-sets. The base of these deposits (i.e. their contact with underlying aeolian facies) is undulose and sharp (Fig. 6). The unit is grey and vague flat lamination laps onto the underlying aeolian dune topography (Fig. 5a) and beds are bounded by flat or current-rippled top surfaces with a thin layer (mm) of very fine material deposited on top (Fig. 5b). The beds are up to tens of centimetres thick and contain rare clasts of consolidated sandstone. Locally, the contact with the underlying aeolian facies is gradational over 10-20 cm, and comprises slightly overturned aeolian cross-bedding.

In Curaco, massive beds with similar texture are observed interbedded with cross-stratified sandstones (0.4 m thick) at the top of the Troncoso Inferior Member succession, overlying both soft-sediment deformed sandstone and aeolian deposits (Fig. 5f). The scale of the trough cross-stratified beds in association with these massive beds is between 3 and 15 cm, and the palaeocurrent directions, both within and between sets, are highly variable. Locally, wave-rippled strata are seen within these sets. The interbedded intervals have a fairly constant thickness of between 1 and 3 m, commonly with a massive base and a low-relief, erosive basal contact. At the top of the Troncoso interval lies a very well cemented rippled unit. This unit occurs across most of the study area as a fairly uniform, 10-30 cm-thick yellow-rusty coloured sandstone (Fig. 5g). Locally, the surface separating the sandstone from the overlying evaporites is wave rippled. This uppermost part of the sandstone succession is fine- to coarse-grained, fairly well sorted and well rounded. Wave-rippled laminae are a few centimetres thick. The massive character results from transport and deposition by high-velocity flows with high fluid content. As the originally dry, aeolian sand became saturated, the water in the pore spaces reduced the cohesive strength, and the mass became liquefied and travelled downslope as liquefied flows (Lowe 1976). The original cross-bedding is destroyed and the vague flat laminations result from shear stresses during movement (Lowe 1976). The presence of this type of lamination and the sharp lower boundary of these deposits provide evidence for sediment mobilization and transportation rather than in situ homogenization. The rippled tops represent the reworking of the bed tops by turbidity currents in the final stages of the flow (Lowe 1976). Vincelette & Chittum (1981) and Benan & Kocurek (2000) proposed that similar massive-rippled facies in reworked aeolian sandstones in the Jurassic Entrada Sandstone (SW USA) were deposited by liquefied flows evolving into turbidites. In addition to simple saturation and collapse of the dune flanks, massive sandstones are, in addition, believed to have formed by collapse and resedimentation resulting from wave erosion of the lower parts of the dunes. Such a process has been invoked to described similar features from flooded aeolian systems on the Colorado Plateau (Eschner & Kocurek 1986; Huntoon & Chan 1987). Subaqueous currents formed the trough cross- and rippledstratification seen in association with the massive

Fig. 6. The Pampa de Tril outcrop (semi-) perpendicular to the palaeowind direction. The lower section is the continuation of the first. The photograph shows part of the outcrop and displays the extensive and sharp contact between the massive-flat-laminated facies and the aeolian dune facies. Weighted lines within the aeolian section are second-order bounding surfaces.

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sandstones in Curaco. The association of massive and rippled sandstones may favour an interpretation of these beds as density current deposits (Ta and Tc of Bouma 1962). However, the high variability in palaeocurrent direction, the limited thickness and the association with wave-rippled sandstone indicate that the crossbedding was produced by wave-generated currents locally reworking the dunes (Clifton 1976). Similar processes occur in the upper part of wave-dominated shoreface systems and indicate that the water depth was less than 10m (Elliott 1986). Wave ripples such as the ones seen in the topmost part of the reworked succession are formed by oscillatory processes forced by waves on the surface of the water. The interaction is typically restricted to water depths that are less than half the wavelength of the waves (Clifton 1976). The extensive cementation at the top of the succession may be associated with the downward percolation of evaporiterich fluids from the overlying deposits.

Sedimentary architecture The Troncoso Inferior Member reaches a maximum thickness of 50 m and underlies the evaporites of the Troncoso Superior Member. Overall, the succession shows a drying-upwards motif. The basal parts of the Troncoso Inferior Member are fluvial and the upper parts aeolian. The contact between the fluvial deposits and the underlying marine deposits of the Agrio Formation and Chorreado Member (Huitrin Formation) is sharp and locally erosive. It is interpreted as a sequence boundary (sensu Posamentier & Vail 1988; Veiga et al 2005). The top of the Troncoso Inferior Member is a sharp contact with the overlying evaporites. This can be classified as both a marine flooding surface (sensu Van Wagoner et al. 1988) and a depositional super-surface (sensu Kocurek & Havholm 1993). In the following section the facies architecture is discussed using data from four 'architectural element analysis' panels from the two key areas (Pampa de Tril and Curaco). Special reference is paid to the preserved aeolian topography and the soft-sediment deformed and reworked interval associated with the flooding of the dune field. Pampa de Tril (two panels) Description. The Pampa de Tril outcrop lies in the northern part of the study area where the Troncoso Inferior Member attains a vertical thickness of up to 50m (Fig. 1). The outcrop

stretches 3.5 km SE-NW, perpendicular to the palaeowind direction (Fig. 6), and 3 km SWNE, parallel to the palaeowind direction (Fig. 7). In the western parts of the Pampa de Tril outcrop the continental deposits lie unconformably on top of ammonite-bearing marine shales of the Agrio Formation. In the eastern parts of the SE-NW panel (Fig. 6) the sandstones lie above a thin (up to 2 m) interval of fine-grained shallow-marine sandstones of the Chorreado Member (Figs 2 & 6). The fluvial package attains a total thickness of 15 m in the NW part of the outcrop, and pinches out to the SE where aeolian facies directly overlie the Chorreado Member. The basal parts of the Troncoso sandstones in Pampa de Tril are composed of erosive and massive-crossbedded fluvial deposits with abundant rip-up clasts. Thick, predominantly massive beds (c. 2 m) pass upwards into thinner beds (tens of centimetres) that are dominantly cross-bedded. The foreset dips within the basal cross-bedded sets vary, but are directed dominantly towards the SE (N130°, n = 244). Horizontal and finely laminated and small-scale cross-bedded aeolian facies occur interbedded with fluvial facies as eroded remnants (western part of the panel in Figs 6 & 7). Remnant and complete aeolian dune bedforms have a height up to 2.5 m. Higher up in the fluvial succession more varying dip directions are dominantly northerly directed (n = 56). Planer-laminated and crossbedded sets are interbedded with fine-grained horizons forming stacked packages. In places accretion co-sets are observed. The fluvial material becomes better sorted and finer grained higher up. Stacked sets of cross-bedded aeolian strata locally overlie the fluvial deposits in the SW (below the panel in Fig. 6) (log 5 in Fig. 3). These sets display small-scale cross-beds that have N-trending foreset directions (N041°, n = 38), and are thin (c. 0.5 m) and laterally discontinuous. The top of this local deposit is cut at its top by a thin, cross-bedded and discontinuous fluvial interval. Across most of the outcrop largescale cross-bedded aeolian strata directly overlie fluvial deposits. The base of the aeolian succession is sharp and comprises granular size grains. Two NNE-SSW-trending ridges dominated by large-scale cross-bedded aeolian deposits occur in the panel in Figure 5. These are 0.5 and 0.8 km wide and 2 km apart. Between the ridges occurs a very thin (less than a few metres) interval of aeolian sandstone. This upper and dominantly aeolian part of the section is up to 35 m thick and comprises cross-bedded sets (up to 7 m thick). The base

Fig. 7. The Pampa de Tril correlation panel parallel to the main palaeowind direction. The areas on the panel with no stratification marked, as well as the base of the panels, are poorly exposed. Third-order bounding surfaces are dashed lines, second-order bounding surfaces are weighted lines.

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of the large-scale cross-bedded interval contains granules. Beds dip dominantly towards the north (N005°, n = 103), but dip directions are locally (in the NW) bimodal towards the north and SW (Figs 5 & 6, respectively). The palaeowind directions at the base of the western ridge are dominantly westerly directed. In the eastern ridge, the palaeowind directions are more variable (Fig. 6). Third-order bounding surfaces are very common within both of the ridges within the large-scale cross-bedded facies (Fig. 7). At the top of the aeolian interval a variety of deposits are observed. In the eastern part of the Pampa de Tril outcrop (Fig. 6), reworked deposits of massive-flat laminated character have a measured thickness up to 15 m. The thickness of the deposits decreases to the west where the interval is completely absent. Similar infilling character of the massive-flat-laminated reworked deposits is seen in the SW parts of the Pampa de Tril outcrop (Fig. 7). Here thinner ( l - 2 m thick) and locally developed beds fill areas of generally lower-dune-topography. The lower contact with the underlying aeolian cross-beds is both sharp (erosional) as well as gradational. The gradational contact is 1020 cm thick and comprises slightly overturned aeolian cross-bedding. The massive-flat-laminated beds in this area extend 50 m laterally and the flat lamination is a few centimetres thick (Fig. 5b). In the SW parts of the outcrop (Fig. 7), rare, small water-escape structures (scale of a few centimetres) are seen at the base of the reworked beds, and rare clasts of lithified sandstone are observed. In areas where no massive-flat-laminated deposits are observed, a 0.1-0.2 m-thick wellcemented wave-rippled sandstone unit sits at the top of the succession (logs 3 and 4 in Fig. 3). Locally, there are no signs of reworking or soft-sediment deformation. Interpretation. Continental sedimentation began with fluvial systems that eroded into the underlying marine Agrio Formation and Chorreado Member of the Huitrm Formation. The erosion incorporated mud and carbonate clasts into the fluvial deposits. Higher in the fluvial succession, planar-laminated and small-scale cross-bedded sets interpreted to have been formed by sheet floods are interbedded with aeolian sandsheet and dune deposits. Individual event beds are stacked into multistorey sets separated by thin muddy horizons that represent waning flow and testify to the ephemeral nature of the depositional events. Evidence for a more confined, semi-permanent flow is seen

towards the middle of the section where downstream and laterally accreting bar complexes are observed. Local fluvial incision is seen in the SW parts of the outcrop (Fig. 6), where a thin cross-bedded fluvial bed overlies stacked sets of small-scale cross-bedded aeolian sandstone. The upwards transition from fluvial to aeolian accumulation implies that the climate became dryer. The wind reworked the underlying fluvial deposits into wind-rippled sandsheets and small aeolian dunes. The initial aeolian dunes migrated towards the NE. Successively, dune cross-sets became larger, and large-scale cross-bedded sets dominate the upper aeolian section of the Troncoso Inferior Member. The granular base of this main aeolian unit indicates a period of deflation of fluvial material before aeolian sedimentation started. Aeolian cross-sets (up to 7 m) build up a total thickness of 35 m. The large-scale crossbedded sets are separated by second- and third-order bounding surfaces (dominantly third order) that are more widely spaced higher up in the succession (Fig. 3), and shows that the active wind regime had stabilized by this time. The bimodal dip pattern of the cross-beds seen in the western ridge in Figure 6 and the palaeowinds measured in the panel in Figure 7 indicate that the dunes were linear. The linear dunes carried stacked sets of sinuous elements that migrated along the dunes, as described from the modern Namib Sand Sea by Bristow et al (2000). The dunes were elongated in a NNESSW direction and had in outcrop a 2 km wavelength. The cross-beds at the base of the western ridge also imply that the dunes were migrating mainly towards the NW at the onset of sedimentation (Fig. 6) (similar to examples by Rubin & Hunter 1985; Bristow et al. 2000), and subsequently developed into more stabilized linear dunes. The dunes were of more complex type in the eastern parts of the Pampa de Tril outcrop. Between the two linear dune ridges in Figure 5, thin aeolian deposits, probably sandsheet deposits and small undeveloped dunes, occupy the interdune area.

The flooding When the Troncoso dune field was flooded by the Troncoso Sea the dunes became saturated with water and the upper parts were subsequently liquefied. This induced mass-flow processes that created a massive-flat-laminated unit which accumulated within topographically low areas. The flat lamination developed as traction carpets in high-density and sandy turbidity

THE TRANSGRESSION OF A CRETACEOUS ERG flows, or by shear stresses on the base of liquefied flows. Deposition by liquefied flows is favoured as the massive-flat-laminated deposits are fairly fine grained and have a rippled top surface. Water-escape structures due to liquefaction can be seen locally at the base of these beds. During late stages in the transgression, and prior to the deposition of evaporites, local subaqueous reworking took place and a thin interval of wave ripples developed locally at the top of the succession. No wave-rippled interval was seen on top of massive-flat-laminated facies and this may indicate that the mass flows took place after the succession became subaqueously reworked. The liquefied flows may have incorporated the wave-rippled section (seen as clasts) into the sediment-water mixture that travelled down the dune face. Sections without evidence of reworking, by waves or liquefaction, might indicate rapid drowning to a depth below wave base. Curaco (two panels) Description. Two panels were produced for the area around Curaco (Fig. 1). One 100 m-long panel strikes WNW-ESE and is perpendicular to the dominant palaeowind direction (A, Fig. 8). The second panel is 125 m long and is orientated towards the north (parallel to the palaeowind direction) (B, Fig. 8). The basal part of the Curaco succession is 810 m of fluvial deposits that rest unconformably on the Agrio Formation, similar to the Pampa de Tril sections. In Curaco, most of the fluvial sandstones are cross-bedded and flat laminated, and are better sorted than those in Pampa de Tril. The lower sets are trough cross-bedded (tens of centimetres thick) and have a principal palaeocurrent direction towards the ESE (N110°, n = 31). Higher in the fluvial interval, current directions are more variable with a mean trough direction toward N040°. In this part of the succession the fluvial deposits are better sorted and finer grained than in the lower parts. No mixed fluvial and aeolian deposit and no finer grained horizons were observed. The top of the fluvial deposits is well cemented and forms a laterally extensive peneplain surface that is overlain by aeolian deposits. The base of the overlying aeolian succession comprises 0.5 m of horizontal and finely laminated sandstones. Where this interval is exposed, it is overlain by stacked sets of 0.51 m-thick cross-bedded aeolian dunes. These trough to wedge, planar, small-scale crossbedded sets are not laterally continuous (a few metres) and together they form a package up to

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4 m thick (Fig. 8). Within the basal cross-sets, foresets are coarse grained and dip towards the east (N074°, n = 25). The small-scale cross-sets are separated by first-order bounding surfaces that have a lateral extent and are locally overlain by coarse-grained and granular material. Higher up in the aeolian succession large-scale cross-bedded sets (up to 4 m) are separated by second-order bounding surfaces, containing subsequent third-order bounding surfaces. These cross-bedded units make up the upper part of the aeolian succession that reaches a maximum thickness of 35 m. In the correlation panel perpendicular to the dominant wind direction the foresets dip unimodally towards the NE (N057°, rc-479) at high angles (20°-30°) (Fig. 8). The majority of the dip directions lie between N020° and N060°. Some inconsistency is seen within trough-shaped sets (Fig. 8). For example, one trough showed varying dip directions from N330° to N064°. The troughs are up to 70 m wide and have low curvature. The bases of the sets are sharp second-order bounding surfaces that are locally overlain by wind-rippled and fine-laminated strata. Towards the top of the succession large-scale cross-bedded aeolian sets are bounded by planar second-order bounding surfaces (B, Fig. 8). In the correlation panel B, parallel to the wind direction, subhorizontal bounding surfaces separate uniformly dipping lee-face deposits. Numerous third-order bounding surfaces are seen within the larger cross-sets. In the Curaco sections, the uppermost parts of the studied succession are different to Pampa de Tril, and are dominated by soft-sediment deformed and wave-reworked deposits. The convolutedly folded soft-sediment deformed units are locally interbedded with dish structures (Fig. 8A) and are observed where the aeolian topography is slightly higher. These intervals are less than 2 m thick and extend for about 40 m with varying thickness and relative proportion of the two different types of deposits (Fig. 8). Both types of deposits are separated from the underlying aeolian dune facies by a thin massive and/or convolutedly folded unit in the form of simply overturned aeolian laminae. Laterally, the soft-sediment deformed units die out or pass into interbedded massive-flatlaminated and cross-stratified units. The softsediment deformed deposits in the Curaco succession are almost exclusively overlain by a massive, a cross-stratified or an interbedded interval of the reworked massive (to flat-laminated) and cross-stratified facies. The massive-flat-laminated interval extends for 100 m with a fairly constant thickness of a

Fig. 8. The Curaco Canyon correlation panels. A is striking perpendicular to the dominant dip direction. B is perpendicular to A and is parallel to the palaeowind direction. Different types of soft-sediment deformation and reworking are seen at the top of the succession and more detailed drawings are highlighted.

THE TRANSGRESSION OF A CRETACEOUS ERG few tens of centimetres and occurs generally in association with reworked cross-stratified beds. Locally, they are interbedded to form intervals l-3m thick (Fig. 8A). The nature of the reworked cross-stratified strata varies across the Curaco area, where it locally forms stacked sets up to a total thickness of 4 m (Fig. 8). In panel A, Figure 8, cross-stratified unit a few tens of centimetres thick, forms a laterally continuous interval at the top of the succession extending for approximately 100 m laterally. The 0.1-0.2 m-thick well-cemented waverippled (locally cross-laminated) unit observed at the top of the section in Pampa de Tril is also present at Curaco. In Curaco, the waverippled unit covers most of the area with the exception of areas with relatively low preserved aeolian topography (below c. 25 m). There are also areas with no signs of reworking or softsediment deformation and the evaporites lie directly on top of aeolian deposits (panel A, Fig. 8). Interpretation. A SE-flowing fluvial system produced the first continental deposits in Curaco. The lack of muddy intervals between cross-stratified sets and the lack of interbedded aeolian deposits indicate that these fluvial systems were less ephemeral than those at Pampa de Tril. A sandsheet was the first aeolian sediment deposited in Curaco. The earliest dune deposits were not laterally extensive, which suggests that the dunes were small and that the system was undersaturated with sediment. The abundance of scouring and the coarseness of the initial dunes also favour an interpretation of a system with a restricted supply of dry and fine loose sand. Low sand input may reflect a wet climate (and substrate) or proximity to fluvial systems (Kocurek & Nielson 1986), both of which would trap sediment. The thin windrippled and granular horizons between the sets indicate that the area was periodically flooded. Wind winnowed the finer grained material leaving a granular lag. As the climate became dryer, finer sand was brought into the area and the dunes became larger. Dune cross-sets (up to 5 m) built up dunes at least 35 m high. First- and third-order bounding surfaces that separate cross-sets resulted from dune migration and changes in wind direction, respectively. The foresets dip with a uniform geometry towards the NE (N057°), have a fairly steep angle and there is a low abundance of wind-ripple strata; this evidence implies that the dunes were barchanoidtransverse ridges. The troughs observed at

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the top of the section are perpendicular to the dominant wind direction and indicate that the dunes in the area were sinuously crested. The width of the troughs were at least 60-70 m, which is the preserved trough width recorded in outcrop. The reworking and soft-sediment deformation in Curaco were fairly complex and involved a series of events. Initial soft-sediment deformation caused by liquefaction and rearrangement of the original dune cross-bedding created a convolutedly folded deposit at the time of liquefaction followed by upwards water escape that formed the dish structures. Later subaqueous processes partially eroded the aeolian deposits and the locally present soft-sediment deformed deposits. Waves cut into the underlying, slightly modified, dune topography and induced liquefied flows that created massive sandstones. These massive sandstones are locally interbedded with a cross-stratified facies. The cross-stratification resulted from shallow-marine reworking. The final stage of reworking was the deposition of a regionally extensive wave-rippled unit, formed as the water depth continued to increase. The absence of the rippled unit in palaeotopographic lows implies that these areas were below the wave base level. Given that the maximum preserved topography is approximately 35 m it can be concluded that the waves were small. This is consistent with the basin being relatively sheltered (Legarreta & Uliana 1991). Underneath the reworked deposits that form laterally extensive and fairly thin units (c. 23 m in total) the preserved aeolian topography has low relief. This might be caused either by the reworking of material from the dune tops, which may have been transported elsewhere, or because the dune setting was of fairly constant height across the area. The latter interpretation is favoured, as there are still signs of preserved dune topography. Further evidence is the lack of marine strata and reworked material in areas of lower relief within the dune topography. The top parts of the dunes would have been extensively reworked if the wave energy was high enough and the reworked material accumulated in areas of lower relief. However, areas that lack evidence for reworking and soft-sediment deformation may have been sheltered from high-energy wave action. Depositional model - discussion The data presented above provide information on the Troncoso Inferior Member depositional system prior to the Late Aptian flooding of the

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Fig. 9. The depositional model of the fluvial and aeolian sedimentation prior to the flooding. The sedimentation in the eastern parts of the area is constrained only by subsurface data.

basin. As well as having important implications for our understanding of the stratigraphy of the Neuquen Basin and the evolution of the Andes, the data have other, more generic, applications. Flooded dune systems are important because, under normal conditions of bedform climb, only a small proportion of the dune is preserved. When dune fields are flooded the depositional system becomes fossilized. This provides a crucial insight into the relationship between internal architecture and dune morphology (Mountney et al 1999). The data presented above also reveal aspects of the processes and products that occur when dune systems are flooded. This has important implications for understanding other flooded aeolian systems where there are less data available, such as the Permian Weissliegend of the southern North Sea Basin (Glennie & Buller 1983; Stromback & Ho well 2002). In the following discussion we shall consider the Troncoso depositional system both prior to, and after, the transgression. Before the flooding As the shelf became exposed by sea-level fall in the Early-Late Aptian a fluvial system came to occupy the whole study area (see discussion in Veiga et al. 2005). Early rivers were erosive but they developed into a partly ephemeral braided system with the main transport direction towards the north (Fig. 9). In Curaco, deposition started off as an extensive aeolian sandsheet and small aeolian dunes. The sediment supply was low, and the building material for the dunes coarse and reworked from the underlying fluvial strata. As the climate became more arid, finer grained material was brought into the basin and larger dunes could develop. In the northern parts of the study area (Pampa de Tril)

linear type dunes with heights up to 35 m and 2 km wavelength were created. In Curaco, smaller barchanoid ridges developed (Fig. 9). The wavelength between the linear dune ridges seen in Pampa de Tril is similar to modern examples seen in the Rub' al Khali Desert (Saudi Arabia) where the mean wavelength is 2.18 km, and the SW Sahara (Mauritania) where the mean wavelength is 1.93km (McKee 1979). The mean width of linear dunes in these areas is 1.21 and 0.94 km, respectively. This is wider than the linear dunes seen in Pampa de Tril where the apparent width is 0.5-0.8 km. The Troncoso dunes in this study may have been smaller and less developed, but the bimodal cross-beds support an interpretation of them as linear dunes. Evidence from Besler (1975), Rubin & Hunter (1985) and Bristow et al. (2000) shows that linear dunes deposit transverse-type strata if these migrate laterally, i.e. if one wind direction is more dominant than another. This is similar to the cross-beds observed at the base of the linear dunes seen in Pampa de Tril that comprise transverse-type strata. Although Bristow et al. (2000) note that linear dunes may deposit abundant trough cross-bedded strata, in Curaco the lack of bimodal dip directions, as well as the highly regular geometry of the dune foresets (parallel to the palaeowind direction), implies that the dunes were barchanoid-trans verse in this area. The 35m height is similar to many modern transverse and barchanoid ridges (Lancaster 1983; Kar 1990). The change from transverse to linear dunes (the distance between Pampa de Tril and Curaco) occurs over a similar distance (25 km in distance) to changes of the same dune types in parts of the modern Namib Desert (Lancaster 1983). When comparing the Troncoso and the Namib examples, the linear

THE TRANSGRESSION OF A CRETACEOUS ERG dunes in some parts of the Namib Desert (adjacent to transverse dunes) are more compound, but the spacing and height of the dunes are similar to the Troncoso. The interdune areas observed in Pampa de Tril are starved of sediment. Evidence like this may suggest that proximity to the palaeoshoreline (towards the north) may have influenced the sedimentary pattern. In Curaco aeolian dune sedimentation seems to have been more persistent as no extensive interdunes are observed. The change in bedform type, from transverse to linear, might have resulted from a relative decrease in sand budget and an increase in wind variability downwind of the transverse bedforms (similar to the southern North Sea: George & Berry 1993). George & Berry (1993) refer to linear dunes as sand-passing bedforms, and barchans and transverse dunes as sand-moving bedforms. A varying wind regime would have reduced the sediment-carrying capacity and the sand would be trapped in the upwind, and more developed, dune field (in this study, the Curaco area). The apparently wetter environment that existed in the northern parts of the study area may also cause the decrease in sand. Factors like these would have stabilized the substrate, and the grain transport would have been hindered as the grains adhered to the damper surface. Aeolian sedimentation ended abruptly in the Late Aptian when the sea inundated the Troncoso dune field. The transgression came from the NW through a narrow embayment of the Pacific realm. The initial flooding drowned the dunes but water depths were shallow enough to allow wave reworking in addition to soft-sediment deformation (see discussion in Veiga et al. 2005). The outcrops provide excellent evidence for understanding the internal aeolian architecture of dunes, and especially the poorly understood linear dunes. The linear dunes observed in Pampa de Tril show evidence for bimodally dipping foresets, but also provides evidence that the initial sedimentation of these ridges commenced with dunes depositing 'transverse-type' strata. Unimodally dipping strata is generally seen as evidence for a lateral sediment transport direction. This type of evidence from the rock record suggests that the preservation of linear dunes may have been misinterpreted as transverse types in previous studies, as normally only the lower parts of the dunes are preserved. Laterally migrating linear dunes have also been suggested in other studies, for example by Rubin & Hunter (1985) and Bristow et al (2000).

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Post-flooding A range of different soft-sediment deformation and reworking processes modified the dune topography of the Troncoso dune field at the time of, and after, the transgression (Fig. 10). These processes are reflected in the facies deposited at the top of the uppermost aeolian unit. The dune field was quietly submerged when the sea level rose in the saline basin centre to the north of the area. The water soaked into the dune sand because of its high porosity and the dunes became unstable as pressure changed internally. The pore spaces within the aeolian dune deposits expanded as they filled with water, and the sand grains subsequently became dispersed within a sand-water mixture. The dunes were liquefied and water escape and mass flows were triggered. In Curaco, early water- and air-escape processes modified the aeolian dune topography and formed convolutedly folded and dish structures (Fig. 8). Although it is proposed that this in situ soft-sediment deformation was associated with the flooding, it is important to consider whether it may have occurred before the sea inundated the basin. Soft-sediment deformation has been observed in many aeolian systems and is attributed to seismic events, cyclic loading by storm waves or changes in water-table level. Seismicity is discounted as a driving mechanism as there is no evidence of convolution on a regional scale (the facies were not seen in Pampa de Tril). Changes in water-table level may have formed the convolutedly folded and dish structures, but as these are not seen elsewhere in the Troncoso system they are also discounted, unless they were related to the transgression. Storm waves may have formed these types of facies, as waves may have been active both at the onset of the transgression and after the dunes were drowned below the storm wave base. As the contact between the soft-sediment deformed deposits and the overlying massive and cross-stratified facies is erosional, the formation of the soft-sediment deformed strata is interpreted to have been before shallow-marine reworking (Fig. 10). Although marine processes reworked the dunes, no large volumes of marine strata were deposited. Only a maximum of 3 m of cross-stratified, massive and waverippled deposits are seen. More marine sediment is seen in the Curaco area compared with the Pampa de Tril and was preserved after deposition due to an increase in the water depth during the flooding. This may suggest that the water depth in the area around Curaco was shallower than

Fig. 10. Proposed model for the sequence of processes that formed the flood-related facies after the Troncoso dunes were flooded.

THE TRANSGRESSION OF A CRETACEOUS ERG around Pampa de Tril. The amount of reworking may also have depended on the time available for marine processes to redistribute material. Oscillatory forces reworked the top of the sandstone succession to form a thin wave-rippled interval in areas with higher preserved dune topography. The deposition of the wave-rippled interval must have occurred in fairly quiet water as it simply draped the top parts of the preserved topography. The local absence of a wave-rippled interval in both Pampa de Tril and Curaco was in areas of generally lower topography and may imply that the water depth was too great in those areas, i.e. the dunes drowned quickly to a depth below the fair-weather wave base. The reworked massive-flat-laminated unit is believed to have been formed by liquefied flows in an already drowned system. The liquefied flows deposited the most voluminous units of all flood-related processes (up to 15 m thick). However, it is interesting to consider whether these flows occurred before or after the marine reworking. Assuming that the topmost wave-rippled interval accumulated on all the dune highs across most of the study area, its absence may indicate that the liquefied flows reworked and removed it. A massive-flat-laminated facies has also been observed by Benan & Kocurek (2000) in the Jurassic Entrada Sandstone in New Mexico where this facies filled, and on-lapped, the underlying dune topography, similar to the examples seen in the Pampa de Tril outcrops. Benan & Kocurek (2000) also interpret this type of facies as having been deposited by mass-flow processes. This geometry (infilling) was also proposed for the southern North Sea (UK) where the Weissliegend facies fills in the preserved topography of the Rotliegend dunes (Stromback & Howell 2002). In comparison, the marine reworked strata are seen as thin (0.1-3 m thick) laterally extensive units. High preservation of dune topography implies that the transgression of the Troncoso dune field was fairly rapid but of fairly low energy. If the transgression was slower, there would have been more time for marine reworking, which would have permitted less dune preservation but more voluminous marine deposits. There would also be evidence of marine reworking across the whole area. If the transgression was slow and/or of higher energy no dune topography would have been preserved. Conclusions The outcrops of the Troncoso Member constitute a great opportunity to study flooded aeolian systems as a large amount of the dune

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topography was preserved after being transgressed. The three-dimensional architecture allowed study of the preserved dunes together with the processes that reworked and caused soft-sediment deformation of them during the flooding. The reworking and soft-sediment deformation could further be related to the amount of dune preservation. The transgression caused extensive reworking and soft-sediment deformation of the dunes, as is demonstrated by the various flood-related facies at the top of the sandstone succession (Fig. 10). In the south (Curaco), towards the basin margin, the flooding initiated liquefaction followed by water escape. Water escape produced convolutedly folded bedding and dish structures within slightly higher preserved dune topography. The modified dune topography in Curaco was further reworked by waves to form marine strata. When the waves undercut the dunes they triggered local liquefied flows that created massive sandstone. A thin wave-rippled interval is present across most of the area, but absent in areas with massive and flat-laminated facies. This implies that the liquefied flows may have eroded the wave-rippled deposits or that these were not deposited at all (i.e. were deposited below the wave base). Liquefied flow deposits, in the form of massive-flat-laminated facies, were seen only in the Pampa de Tril area. Voluminous reworking of the linear dunes by liquefied flows deposited massiveflat-laminated facies that came to occupy lower areas within the dune topography. The northern parts of the field area (Pampa de Tril) were quickly submerged and there would have been no time to develop a wave-rippled facies. Some areas show no signs of reworking or softsediment deformation. The massive and flat-laminated facies deposited by liquefied flows forms the greatest volume of reworked strata, and may imply that these deposits have greatest significance when studying flooded dune systems in the subsurface. Different types of marine facies were deposited as fairly thin (c. 2-3 m) laterally extensive units. The flood-related facies do not seem to influence the amount of dune preservation. Instead there is a different amount of preserved aeolian relief of the strata below the flood-related deposits. In the areas with more liquefied flows, the relief is greater compared with the dunes that were reworked by marine processes, although this might also be due an artefact of the dune setting at the time of the transgression. The amount of marine strata, as well as the large amount of preserved topography, indicate

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that the flooding of the Troncoso dune field was rapid but of fairly low energy. Below the flood-related units at the top of the succession, the Troncoso Inferior Member was deposited as a sand-rich facies in both fluvial and aeolian environments. The fluvial deposits are dominated by cross-bedded and plane-laminated sandstones. The geometry of these sandstones indicates that they were deposited within low-sinuosity fluvial channels and as sheet floods. When the climate became drier large aeolian dunes started to develop and fluvial deposition was restricted to interdune corridors. Evidence that an extensive aeolian dune field began to accumulate after the main fluvial phase is present across the whole study area. Aeolian sets were preserved as dunes/draas, at least 35 m high. The sedimentation style changed downwind (northwards) from trans verse/barchanoid into linear dunes. The different types of dunes accumulated within 25 km from each other, with dune scale and geometry similar to parts of the modern Namib Desert (Lancaster 1983). We would like to thank ARCO British Ltd (BP Amoco) for funding this research. Further we would like to thank K.W. Glennie and J.D. Marshall for their constructive feedback on this paper as a part of a thesis undertaken by A. Stromback at the University of Liverpool, UK ('An evaluation of the Weissliegend facies of the UK, Southern North Sea', 2001). R. Blakey, D. Loope and A. Tripaldi are thanked for the review of the paper in its present form.

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CLIFTON, H.E. 1976. Wave-formed sedimentary structures - a conceptual model. In: DAVIES, R.A., JR & ETHINGTON, R.L. (eds) Beach and Nearshore Sedimentation. SEPM, Special Publications, 24, 126-148. COMERON, R.E. 1990. Trampas estratigraficas en sedimentos de origen eolico (Un estudio particularizado del Mb. Troncoso Inferior en el area de Chihuido de la Sierra Negra, Provincia de Neuquen). Boletin de Informaciones Petroleras, 1990, 2-7. ELLIOTT, T.E. 1986. Siliciclastic shorelines. In: READING, H.G. (ed.) Sedimentary Environments and Facies, 2nd edn. Blackwell Science, Oxford, 155-188. ESCHNER, T.B. & KOCUREK, G. 1986. Marine destruction of eolian sand seas: origin of mass flows. Journal of Sedimentary Petrology, 56, 401-411. ESHNER, T.B. & KOCUREK, G. 1988. Origin of relief along contacts between eolian sandstones and overlying marine strata. AAPG Bulletin, 72, 932-943. FRANZESE, J.R. & SPALLETTI, L.A. 2001. Late Triassic-early Jurassic continental extension in southwestern Gondwana: tectonic segmentation and pre-break-up rifting. Journal of South American Earth Sciences, 14, 257-270. GEORGE, G.T. & BERRY, J.K. 1993. A new palaeogeographic and depositional model for the Upper Rotliegend of the UK Sector of the Southern North Sea. In: NORTH, C.P. & PROSSER, D.J. (eds) Characterization of Fluvial and Aeolian Reservoirs. Geological Society, London, Special Publications, 73, 291-319. GLENNIE, K.W. & BULLER, A.T. 1983. The Permian Weissliegendes of N.W. Europe: the partial deformation of eolian dune sands caused by the Zechstein transgression. Sedimentary Geology, 35,43-81. GULISANO, C.A. & GUTIERREZ PLEIMING, A.R. 199 Field Guide - The Jurassic of the Neuquen Basin. Mendoza Province, Buenos Aires. Secretaria de Minera de la Nacion. Direction Nacional Servicio Geologico Publication, 158. HOWELL, J.A. & MOUNTNEY, N. 1997. Climatic cyclicity and accommodation space in arid to semi-arid depositional systems: an example from the Rotliegend Group of the UK southern North Sea. In: ZIEGLER K., TURNER, P. & SAINES, S.R. (eds) Petroleum Geology of the Southern North Sea: Future Potential. Geological Society, London, Special Publications, 123, 63-86. HUNTER, R.E. 1977. Basic types of stratification in small eolian dunes. Sedimentology, 24, 361-387. HUNTOON, J.E. & CHAN, M.A. 1987. Marine origin of paleotopographic relief on eolian White Rim Sandstone (Permian), Elaterite Basin, Utah. AAPG Bulletin, 71, 1035-1045. KAR, A. 1990. Megabarchanoids of the Thar: their environment, morphology and relationship with longitudinal dunes. Geographical Journal, 156, 51-61. KOCUREK, G. 1981. Significance of interdune deposits and bounding surfaces in aeolian dune sands. Sedimentology, 28, 753-780.

THE TRANSGRESSION OF A CRETACEOUS ERG KOCUREK, G. 1996. Desert aeolian systems. In: READING, H.G. (ed.) Sedimentary Environments: Processes, Fades and Stratigraphy, 3rd edn., Blackwell Science, Oxford, 125-153. KOCUREK, G. & HAVHOLM, K.G. 1993. Eolian sequence stratigraphy - a conceptual framework. In: WEIMER, P. & POSAMENTIER, H.W. (eds) Siliciclastic Sequence Stratigraphy. AAPG Memoirs, 58, 393-409. KOCUREK, G. & LANCASTER, N. 1999. Aeolian system sediment state: theory and Mojave Desert Kelso dune field example. Sedimentology, 46, 505-515. KOCUREK, G. & NIELSON, J. 1986. Conditions favourable for the formation of warm-climate aeolian sand sheets. Sedimentology, 33, 795-816. LANCASTER, N. 1983. Controls on dune morphology in the Namib sand sea. In: BROOKFIELD, M.E. & AHLBRANDT, T.S. (eds) Eolian Sediments and Processes. Developments in Sedimentology, 38., Elsevier, Oxford, 521-541. LEGARRETA, L. 1985. Andlisis estratigrafico de la F. Huitrin (Cretacico Inferior), provincia de Mendoza. Doctoral Thesis, Universidad de Buenos Aires. LEGARRETA, L. & GULISANO, C.A. 1989. Analisis estratigrafico secuencial de la Cuenca Neuquina (Triasico Superior-Terciario Inferior), Argentina. In: CHEBLI, G.A. & SPALLETTI, L.A. (eds) Cuencas Sedimentarias Argentinas. Serie de Correlacion Geologica, 6, 221-243. LEGARRETA, L. & ULIANA, M.A. 1991. JurassicCretaceous marine oscillations and geometry of back-arc basin fill, central Argentine Andes. In: MACDONALD, D.I.M. (ed.) Sedimentation, Tectonics and Eustacy — Sea—Level Changes at Active Margins. International Association of Sedimentology, Special Publications 12, 429-450. LOWE, D.R. 1975. Water escape structures in coarsegrained sediments. Sedimentology, 22, 157-204. LOWE, D.R. 1976. Subaqueous liquefied and fluidized sediment flows and their deposits. Sedimentology, 23, 285-308. LOWE, D.R. & LoPiccoLO, R.D. 1974. The characteristics and origins of dish and pillar structures. Journal of Sedimentary Petrology, 44, 484-501. McKEE, E.D. 1979. A Study of Global Sand Seas, US Geological Survey, Professional Paper, 1052, 399407. MIALL, A.D. 1988. Architectural elements and bounding surfaces in fluvial deposits: anatomy of the Kayenta Formation (lower Jurassic), southwest Colorado. Sedimentary Geology, 55, 233-262. MOUNTNEY, N.P., HOWELL, J.A., FLINT, S.S. &

JERRAM, D.A. 1999. Climate, sediment supply and tectonics as controls on the deposition and preservation of the aeolian-fluvial Etjo Sandstone Formation, Namibia. Journal of the Geological Society, London, 156, Ill-Ill.

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POSAMENTIER, H.W. & VAIL, P.R. 1988. Eustatic controls on clastic deposition II - sequence and systems tract models. In: WILGUS, C.K., HASTINGS, B.S., KENDALL, C.G.ST.C, POSAMENTIER, H.W., Ross, C.A. & VAN WAGONER, J.C. (eds) Sea-level Changes: An Integrated Approach. SEPM, Special Publications, 42, 39-45. RAUTMAN, C.A. & DOTT, R.H., JR. 1977. Dish structures formed by fluid escape in Jurassic shallow marine sandstones. Journal of Sedimentary Petrology, 47, 101-106. RUBIN, D.M. & HUNTER, R.E. 1985. Why deposits of longitudinal dunes are rarely recognised in the geologic record. Sedimentology, 32, 147-157. STROMBACK, A.C. & HOWELL, J.A. 2002. Predicting distribution of remoblized aeolian facies using sub-surface data: the Weissliegend of the UK Southern North Sea. Petroleum Geoscience, 8, 237-249. ULIANA, M, DELLAPE, D. & PANDO, G. 1975. Distribucion y genesis de las sedimentitas rayosianas (Cretacico Inferior de las provincias del Neuquen y Mendoza, Argentina). // Congreso Iberoamerica de Geologica Economica, I, 151-176. ULIANA, M.A. & LEGARRETA, L. 1993. Hydrocarbons habitat in a Triassic to Cretaceous Sub-Andean setting: Neuquen Basin, Argentina. Journal of Petroleum Geology, 16, 397-420. VAN WAGONER, J.C., POSAMENTIER, H.W., MITCHUM, R.M., VAIL, P.R., SARG, J.F., LOUTIT, T.S. & HARDENBOL, J. 1988. An overview of the fundamentals of sequence stratigraphy and key definitions, In: WILGUS, C.K., HASTINGS, B.S., KENDALL, C.G.ST.C., PASAMENTIER, H.W., Ross, C.A. & VAN WAGONER, J.C. (eds) Sea-level Changes: An Integrated Approach, SEPM, Special Publications, 42, 39-45. VEIGA, R. & Rossi, G. 1992. Analisis Sedimentologico del Miembro Troncoso Inferior (Formacion Huitrin) en el Ambito de la Sierra Reyes, Departamento del Malargiie, Provincia de Mendoza. IV Reunion Argentina de Sedimentologia, 1, 71—78. VEIGA, G.D., HOWELL, J.A. & STROMBACK, A. 2005. Anatomy of a mixed marine-non-marine lowstand wedge in a ramp setting. The record of a Barremian-Aptian complex relative sea-level fall in the central Neuquen Basin, Argentina. In: VEIGA, G.D., SPALLETTI, L.A., HOWELL, J.A. & SCHWARZ, E. (eds) The Neuquen Basin: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 139-162. VINCELETTE, R.R. & CHiTTUM, W.E. 1981. Exploration from oil accumulations in Entrada Sandstone, San Juan Basin, New Mexico. AAPG Bulletin, 65, 2546-2570. WILLIAMS, G.E. 1971. Flood deposits of the sand-bed ephemeral streams of central Australia. Sedimentology, 17, 1-40.

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Relative oxygenation of the Tithonian-Valanginian Vaca Muerta-Chachao formations of the Mendoza Shelf, Neuquen Basin, Argentina P. DOYLE1, D. G POIRE2, L. A. SPALLETTI2, D. PIRRIE3, P. BRENCHLEY4 & S. D. MATHEOS2 Department of Earth Sciences, University College London, Gower Street, London WC1E 6BT, UK Centro de Investigaciones Geologicas, Universidad Nacional de La Plata-Conicet, Calle 1 No. 644, 1900 La Plata, Argentina 3

Camborne School of Mines, University of Exeter, Redruth, Cornwall TR15 3SE, UK 4

Department of Earth Sciences, University of Liverpool, 4 Brownlow Street, Liverpool L69 3GP, UK

Abstract: Organic-rich sediments were deposited in the deeper sectors of the Neuquen Basin during the latest Jurassic and the Early Cretaceous. This paper presents the results of a detailed examination of these deposits in the northern-most extension of the basin, in the Mendoza Province, and explores their wider significance for palaeo-oxygenation studies. The Tithonian-Berriasian Vaca Muerta Formation, the primary source rock for the Neuquen Basin, comprises bituminous shales and interbedded limestones deposited during a major transgression. In the Valanginian, the beginning of a regressive phase enabled the development of shallow-marine carbonates to form the base of the Chachao Formation, which eventually led to extensive biohermal carbonates of the uppermost Chachao Formation. Along the length of the narrow N-S-trending Mendoza Shelf of the Neuquen Basin both units are well exposed, permitting detailed study of the stratigraphy, sedimentology, ichnology and palaeoecology. The analysis of the Tithonian-Valanginian succession in the Salado river valley shows that carbonate production increased up-section. Faunal associations are mostly limited to poorly diverse epibenthos and pseudoplankton in the lower part of the section (Vaca Muerta Formation), with increased diversity in the lower Chachao section, including shallow and deeper infaunal bivalves. A background level of laminated shales to Chondrites bioturbation is typical of anoxic-suboxic conditions. Micritic limestones and carbonate sandstones throughout the section commonly show the development of Thalassinoides suevicus. Relative oxygenation curves based on trace fossils and body fossils were developed and compared. There was a primary substrate control on trace fossil diversity and occurrence, with a primary oxygenation signal provided by body fossil evidence. Interpretation of the palaeo-oxygenation on the basis of trace fossil taxa alone, however, would lead to inaccurate results. This study, therefore, demonstrates the importance of integrated trace and body fossil analysis in the fuller understanding of black shales.

Palaeoenvironmental analysis of oxygendeficient basins has been the subject of many detailed studies over the last 25 years, fuelled in part by the development of trace fossil models of relative palaeo-oxygenation following the initial seminal work by Rhoads & Morse (1971). In particular, Savrda & Bottjer (1986, 1987), Ekdale & Mason (1988), Allison et al (1995) and Savrda (1995) have developed and synthesized models in which trace fossil diversity and burrow diameter are used as a proxy

for relative oxygen levels within specific basins, with a direct relationship usually being recorded between increased diversity patterns/ burrow diameter and increasing levels of dissolved oxygen. This relationship has been tested in a number of ancient sedimentary basins following the development of the original Rhoads & Morse model, but little critical appraisal has been made of it. This is particularly true for the upper, oxygenated conditions, where the occurrence of wide burrow-diameter Thalassinoides

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 185-206. 0305-8719/05/$15.00 © The Geological Society of London 2005.

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and other diagnostic traces have been recognized as important indicators of increased oxygen levels. Other studies have concentrated on analysis of the diversity of body fossil macrobenthos in low oxygen environments (e.g. Duff 1974; Wignall 1990; Etter 1995). Together, body and trace fossil models have provided a useful guide to dissolved oxygen levels in specific basins. Recent studies have shown that the diversity response is more complex and difficult to determine solely with reference to oxygenation (e.g. Sageman & Bina 1997). Certain signals, such as the common development of opportunistic species and the adoption by benthos of pseudoplanktonic habit, are common in black shales (e.g. Wignall & Simms 1990; Doyle & Whitham 1991; Wignall 1993; Etter 1996). Such palaeoecological responses may reflect a primary low oxygen signal, as has been supposed in many studies. However, they may also indicate factors such as environmental instability and the presence of soupy substrates (Wignall 1993; Sageman & Bina 1997) or the capability of mass flow and other environmental factors in introducing oxygen-demanding trace makers in to otherwise anoxic basins (e.g. Follmi & Grimm 1990). The present study describes for the first time the palaeoenvironments of an important oxygen-deficient stratigraphic interval, the Vaca Muerta-lowermost Chachao formations of Tithonian-Valanginian age, in the Mendoza Shelf of the Neuquen Basin (Mendoza Province of Argentina) (Fig. 1). The Vaca Muerta Formation in particular is considered to be the primary source rock for the important Neuquen oilfield in Argentina, and yet it has not, to date, received extensive study in this field. Previous works on the palaeontology and stratigraphy of the Vaca Muerta Formation have been confined to the stratigraphy, aspects of particular macroinvertebrate and vertebrate fossil groups, and details of the sedimentology (e.g. Leanza et al 1977; Leanza 1981; Leanza & Zeiss 1990; Gasparini et al. 1997, 1999, 2002; Spalletti et al. 1999). The Chachao Formation, and particularly its oxygen-deficient lower part, has received similarly scant attention (Mombru et al. 1978; Legarreta et al. 1981; Legarreta & Kozlowski 1981). The aim of this paper is to describe in detail the palaeoenvironments of the Vaca Muerta and lower Chachao formations from an examination of trace fossils and benthic macropalaeontology, and to compare the oxygenation signals provided by both data sources in order to test the veracity of the trace fossil oxygenation model.

Late Jurassic-Early Cretaceous palaeogeography, stratigraphy and sequence stratigraphy of the Neuquen Basin Two main depositional areas can be recognized in the Neuquen Basin (Fig. 1): the Neuquen Embayment (Bracaccini 1970) to the south, and the narrow Mendoza Shelf (where the studied section is located) to the north. The Neuquen Embayment was a wide subcircular bay or gulf that developed behind the Andean magmatic arc. During Jurassic and Early Cretaceous times this region was limited by emergent areas to the west, south and east; from all these points and to the interior of the depocentre, smooth depositional ramp topographies converged. The deepest area of the basin was limited to an elongated strip whose greater extent was oriented NW-SE, oblique to the Andean magmatic arc, which extended north-south (Spalletti et al. 2000). This obliquity between the arc and the axis of the depocentre favoured the development of the Mendoza shelf to the north (Fig. 1). As shown previously by Mombru et al (1978), Legarreta et al. (1993) and Spalletti et al. (2000), this area was characterized by a significant reduction of the amplitude of the basin, which in latitudes lower than 34° S was limited to a thin strip located in the surroundings of the Jurassic-Cretaceous magmatic arc. In most of the Neuquen Basin, the Upper Jurassic-Lower Cretaceous record is characterized by the dark bituminous shales and marls of the Vaca Muerta Formation (Weaver 1931). This unit concordantly overlies the clastic and continental deposits of the Tordillo Formation. The Tordillo-Vaca Muerta contact (Early Tithonian) is an isochronous surface throughout the basin and marks the begining of the marine Tithonian transgression (Leanza 1981). The top of the Vaca Muerta Formation is diachronous (Leanza 1973, 1981; Leanza & Hugo 1977; Leanza et al. 1977) and progradational (Gulisano et al. 1984; Mitchum & Uliana 1985; Legarreta & Gulisano 1989). It includes younger stages when passing from the southern (Middle Tithonian) to the central sector of the basin (Berriasian-Valanginian). In the Mendoza Shelf, the Vaca Muerta Formation is conformably overlain by the marine carbonates and shales of the Valanginian Chachao Formation (Fig. 2). The stratigraphy of these two units has been summarized by Riccardi (1983, 1988), Legarreta et al. (1993) and Gulisano & Gutierrez Pleimling (1994).

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Fig. 1. (A) Location map; (B) palaeogeographic and tectonic sketch; and (C) simplified geological map (modified from Gulisano & Gutierrez Pleimling 1994) of the study area.

The progradational record of the Tithonian to the Berriasian-Valanginian in all of the Neuquen Basin was defined by Legarreta & Gulisano (1989) as the Lower Mendoza Mesosequence (Fig. 2). Legarreta & Gulisano (1989), Legarreta & Uliana (1991, 1996) and Legarreta

et al (1993) have interpreted it as the result of a second-order eustatic cycle, combined with persistent regional subsidence regulated by cooling and thermal contraction phenomena, and associated with a very reduced clastic sediment supply and conditions suitable for the

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Fig. 2. Stratigraphic chart of the Tithonian-Valanginian interval in the Mendoza Shelf. Biostratigraphy from Leanza (1973, 1981), Leanza & Hugo (1977), Leanza et al (1977), Leanza & Zeiss (1990) and Aguirre-Urreta & Rawson (1998).

anaerobic-dysaerobic condensed sedimentation. However, as stated by Hallam (1991) and Spalletti et al. (2000), the regional tectonism, related to the activity of wrench fault systems and the growth of the Andean magmatic arc, cannot be disregarded as an important control in the development of this large cycle. Sequence stratigraphical studies of the Lower Mendoza cycle (Gulisano et al. 1984; Mitchum & Uliana 1985; Legarreta & Gulisano 1989; Legarreta & Uliana 1991; Gulisano & Gutierrez Pleimling 1994; Spalletti et al. 2000) have permitted the recognition of several higher order shallo wing-up wards depositional sequences. Based on sequence thickness and estimated geochronological duration, these authors have

suggested a correlation between the geometry and the arrangement of the observed sequences and the global chart of third-order eustatic variation (Haq et al 1987). The depositional sequences of the Mendoza Shelf are different to those of the Neuquen Embayment; however, at the basin depocentre their identification is quite difficult because of the lack of major unconformities and the uniformity of the sedimentary record (Fig. 2).

Sedimentology Field work was carried out in the central part of the Mendoza Shelf, along the Salado river valley (35°12'S and 69°46'W), where the

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Tithonian-Valanginian succession (170m) is part of a large anticlinal structure oriented N-S and accompanied by several N-S-oriented thrust faults (Fig. 1). The studied section has been measured along the northern and southern margins of the Salado river valley, across the western flank of the anticline. It comprises the upper part of the Vaca Muerta Formation (its lower section is not exposed) and the whole Chachao Formation up to its stratigraphic contact with the overlying Agrio Formation (Figs 1 & 2). Fades analysis The Tithonian-Valanginian succession is dominated by fine-grained siliciclastic facies, nodular and bedded marls, and bioclastic carbonates (Fig. 3). The Vaca Muerta Formation is characterized by fine-grained sediments, mostly black shales, with some dark-grey limy shales and dark marls, formed by suspension fall-out in geochemically anoxic-suboxic environments. Some of these fine-grained sediments are very rich in ammonites and small bivalves. One of the most common features of these deposits, already mentioned by Gasparini et al (1997, 1999), is the small-scale cyclicity consisting of 0.4-2 m alternations of black shales and organic marls (Fig. 3). Nodular limestones (micrites and marls) are associated with the fine-grained siliciclastic rocks; nodule concentrations are rhythmically spaced, and were mostly precipitated on skeletal cores (small to large fragments). Very thin and episodic ash layers are relatively frequent, from 2 to 10 cm thick, occasionally reaching 40 cm thick, and exhibiting sharp bases where the ashes overlie shales. Some massive layers or those with well-developed normal gradation are interpreted as ash-fall deposits and/or volcaniclastic deposits reworked by low-density or dilute gravitational flows. Thin and isolated packstones and wackestones appear within the fine-grained packages. These beds usually exhibit an irregular and/or sharp lower surface and slight normal gradation. They can be either massive, laminated or with symmetrical or asymmetrical combined flow ripples. In the Vaca Muerta Formation, two informal stratigraphical sections can be recognized (Fig. 4). The lowermost one is composed of 75 m of black bituminous shales and (massive and/or nodular) fetid marls and micrites, with subordinated intercalations of coarser grained bioclastic carbonates, as well as thin and fine-grained ash layers. The upper section is composed of a rhythmic succession (23 m) of grey limy shales and wackestones, with isolated intercalations of

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packstone beds. Towards the top of this section, some coarsening- and fining-upwards small cycles (1.5-3 m thick) were identified. The Chachao Formation overlies gradationally the Vaca Muerta Formation, and most of its facies are similar to those of the Vaca Muerta Formation. However, in the Chachao Formation marly beds and bioclastic carbonates are more common. The Chachao Formation shows characteristically a series of wackestones and packstones with abundant molluscan (oyster-rich) shell debris. As discussed below, some of these beds show in situ deep- to shallow-burrowing infaunal bivalves. Other beds are almost exclusively formed of abundant disarticulated and fragmentary bivalve debris. Commonly, these shell beds are composed of a lower interval of deep-infaunal bivalves lying in life position, and a strongly reworked upper bioclastic interval. In the Salado river valley, the Chachao Formation can be subdivided into three informal sections (Figs 3 & 4). The lowermost one (36 m thick) is composed of dark-grey limy shales and dark shales, heterolithic beds (thin alternations of wave-rippled sandstones and mudstones), marls and packstones showing primary structures, such as planar bedding, wave-ripple lamination and hummocky cross-stratification (HCS), and deformational structures (waterescape features, ball-and-pillow). This section shows a well-developed parasequence stacking, composed of cycles averaging 4 m thick. The middle section of the Chachao Formation (Fig. 4) is a 17 m-thick succession characterized by alternating cycles of dark-grey shales and marls, with isolated and thinly bedded, bioturbated, bioclastic wackestones composed of fragmented infaunal shells and commonly with concentrations of the pectinid bivalve Entolium concentrations on top. However, for the most part this section, and especially its middle part, lacks diverse benthonic organisms and more often contains scattered planktonic and pseudoplanktonic remains. However, to the top of the section, marly beds usually show a few infaunal bivalves and scattered semi-infaunal (Pinna) bivalves in life position. The upper part of the Chachao Formation (23 m thick, Fig. 4) is composed of coarsegrained shell beds associated with less frequent marls and dark shales (like those of the middle section). In the lower part of this section, shell beds comprise burrowed grainstones and packstones composed of fragmentary gastropods, some infaunal bivalves, including trigoniids, and oysters. As with the middle section, many of these beds show concentrations of epifaunal

Fig. 3. Lithostratigraphical column for the Tithonian-Valanginian interval in the Salado River section. Sample numbers are on the left of the column.

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Fades and cycles Biogenic carbonate banks, coarse packstones, marls and shales. Well developed parasequences limited by flooding surfaces.

Dark grey limy shales (lower part) and dark shales (upper part) associated with frequent marls and packstones (more common in the lower part). Well developed parasequences limited by flooding surfaces.

Dark grey limy shales, heterolithic intervals, marls and packstones. Cyclic arrangements and some CU sequences (shale->marl->carbonate sandstones). Grey limy shales and wackestones and very isolated packstones. Common shaie-wackestone Cycles. Some CU and FU small cycles at the top.

Black shales and interbedded marls, some of them nodular. Black carbonate-rich shales to the top. Some isolated coarse-grained carbonate beds. Common cycles shale-marl.

Fig. 4. Simplified log of the Vaca Muerta and Chachao formations. Informal sections, facies and cycles. Sample numbers are on the right of the column. CU, coarsening-upwards cycles; FU, fining-upwards cycles.

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bivalves (e.g. Entolium) on top. The upper part of this section is, instead, dominated by autochthonous and para-autochthonous coarsegrained and massive (bioturbated) packstones consisting of amalgamated 0.3-0.6 m-thick tabular layers composed of bored ostreid oyster shells. These organisms are associated with gastropods and epifaunal bivalves (Entolium), as well as infaunal bivalves (trigoniids, Panopea, Protocardium), in life position. In some cases these are found as articulated shells oriented parallel to the bedding planes, even forming a more chaotic and finer grained matrix of bioclastic debris. Environmental interpretation Based on the general geometry of the substrate, Mitchum & Uliana (1985) and Legarreta & Uliana (1991) have proposed a depositional ramp model for the Tithonian-Valanginian units of the Neuquen Basin, therefore devoid of submarine fan or apron deposits as well as large-scale reef structures. For the Neuquen Embayment, Spalletti et al (2000) showed a simple distribution of carbonate and siliciclastic facies associations, and a gradual transition from the shallow-marine areas to the deepest areas without a talus slope or a significant slope break. According to the conceptual model proposed by Burchette & Wright (1992), Spalletti et al. (2000) basinal, outer ramp, middle ramp and inner ramp environments were identified. In the Mendoza Shelf, the general pattern of the Tithonian-Valanginian sedimentation seems to be slightly different. During the Valanginian the ramp geometry of the basin gave rise to a mixed siliciclastic-carbonate platform. The thick bioclastic carbonates of the upper Chachao Formation (Mombru et al. 1978; Legarreta & Kozlowski 1981; Legarreta et al. 1993) suggest the development of an almost continuous build-up (reef structure) along the boundary between the narrow platform to the east and the talus slope to the west and NW. Most of the Vaca Muerta Formation in the Salado section is characterized by fine-grained shaly and carbonate organic-rich sediments, often with a cyclic depositional arrangement. These deposits accumulated in a basinal environment, and represent the suspension sedimentation of siliciclastic and carbonate particles above an anoxic-suboxic seafloor. A shallower, but still basinal, marine environment can be proposed when the dark fine-grained sediments show thin siliciclastic or carbonate sand intercalations. As described in previous papers (Gasparini et al. 1997; Spalletti et al.

1999, 2000) these coarser-grained layers, with evidence of wave reworking and lowdensity tempestite flows, are interpreted as very distal deposits produced by storm processes (Myrow & Southard 1991, 1996; Midtgaard 1996). A depositional setting representing the transition between the offshore and the lower shoreface environments can be proposed for the deposits of the lower Chachao Formation, where the bioclastic carbonate beds are more frequent and the parasequence stacking is clearly developed. In the middle Chachao section, deeper basinal marine conditions, like those of the Vaca Muerta Formation, were re-established. The oyster beds of the Upper Chachao Formation represent a marginal sector of a large 'reef structure, probably developed by the margin of the Mendoza Shelf. The shell beds from the lower parts of the Upper Chachao Formation were formed in shallow and high-energy environments under conditions of reduced rates of terrigenous sediment accumulation (Abbott 1997). They represent a transition zone between shoreface and offshore, with a macrofossil assemblage composed of mixed taxa representing both nearshore and offshore environments. The parasequences composed of autochthonous and parautochthonous oysterrich packstones, marls and dark shales in the uppermost Chachao Formation probably represent a relatively quieter shelf environment. Although oyster beds have been associated with restricted environments, we have also observed the development of oyster banks in the open shallow-marine platform surrounding the Margarita Island (north Venezuela, 11°N latitude). If this was the case in the upper Chachao Formation, the vertical distribution of shell beds can be interpreted in terms of a slightly upwards-deepening succession, from the lower shoreface to the proximal offshore in an open shallow-marine mixed (carbonate-siliciclastic) platform. As such, it is more likely that the low faunal diversity of these beds may be the result of reduced competition by other organisms for space and nutrients (faunally restricted environments) as well as a lack of predation (Glenn & Arthur 1990; Abed & Sadaqah 1998), rather than traditional interpretations of restricted environments. Ichnology The Salado River section exhibits a poorly diverse ichnological assemblage associated with a background of relatively fine-grained substrates (Fig. 5). The transition from the Vaca

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Fig. 5. Trace fossil distribution. 1, Thalassinoides; (1A, type A; IB, type B; 1C, type C; ID, type D; IE, type E; IF, type F); 2, Arenicolites; 3, Chondrites; 4, Gordia; 5, Palaeophycus; 6, Planolites; 7, Phycodes; 8, Teichichnus; 9, Rhizocorallium; 10, small diameter burrows; 11, Gastrochaenolites; 12, Trypanites. A, laminated facies association; B, Chondrites association; C, Thalassinoides association; D, Arenicolites association.

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Muerta to the Chachao Formation is marked by the transition from black mud to carbonate-rich and heterolithic intervals, and therefore the increased possibility of substrate control having a larger impact on the resultant biofacies development. This is particularly true where there are periodic episodes of sand deposition, which contain typically sand-dominated traces such as Arenicolites and Gordia, traces that are otherwise largely absent from the succession. Four principal ichnological associations can be recognized (Fig. 5), which are discussed below. Laminated fades (non-bioturbated association) This facies represents the basic background conditions within the basin. Laminated deposits, lacking bioturbation, are locally interbedded with the Chondrites bioturbated deposits described below. This is a common feature of most black shale successions (Rhoads & Morse 1971; Bromley & Ekdale 1984; Savrda & Bottjer 1986; Wignall 1993) and is clearly related to low dissolved oxygen levels, laminated facies being largely a product of truly anoxic conditions (Wignall 1993). Chondrites association Chondrites is the commonest background-level trace fossil and its trace maker is considered to be the primary bioturbator in homogenizing laminated mud facies such as those in the Vaca Muerta Formation. This is in keeping with the role for this trace fossil as the first to colonize and the last to leave in oxygen-deficient black shale sequences (Bromley & Ekdale 1984). However, thoughout the succession, Chondrites is poorly diverse, with small overall burrow diameters that are on a millimetre scale. This is nothing like the large burrow diameter Chondrites that developed in the oxygen-deficient Lower Jurassic Posidonienschiefer of Germany, where burrow diameters commonly reach up to 5 mm or more (Seilacher 1982a, b). Thalassinoides association Thalassinoides is relatively common and is most frequently associated with the carbonate and heterolithic facies within the Salado section. Thalassinoides is commonly considered as an indicator of oxic conditions, and is therefore an important oxic benchmark in the palaeooxygenation schemes of Savrda & Bottjer (1986). Furthermore, Savrda & Bottjer (1986,

1987, 1989) have suggested that the maximum diameter of the burrows depends on the amount of oxygen in the substrate. The greater the oxygenation the larger the infaunal bioturbation. From this relationship an upwards increase in oxic conditions is interpreted from the increase in the diameter of the Thalassinoides burrows. Thalassinoides is also a determinant of substrate firmness, as the complex galleries of the burrow system are not capable of remaining unsupported in a soupy substrate (Ekdale et al. 1984). Within the Chachao section of the Salado River, at least six types have been recognized, A-F (Fig. 6). These forms are discussed below. Thalassinoides types A-D. Thalassinoides types A-D conform to the ichnospecies Thalassinoides suevicus as currently defined (Ekdale et al. 1984), characterized by regular galleries and straight shafts (Fig. 6). They are developed within the carbonate units in the Salado River section, and the recognition of four separate morphotypes (A-D) is dependent on the differentiation of burrow fill. Thalassinoides type A is characterized by a bioclastic infill that distinguishes it from its otherwise poorly bioclastic host sediment. In this case, the infill is interpreted as the by-product of feeding and predation by the Thalassinoides trace-making organism, and perhaps represents active successive debris chamber infill by predators. Thalassinoides type B displays a sand passive infill with some bioclastic pockets, the bioclasts again probably being derived from predation. Infill here may conform to both passive and active infilling by the trace maker, and, dependent on the burrow, may be interpreted as a function of activities in the range from both soft to firm grounds. Thalassinoides type C has a passive sand infill alone, with the sand filling the empty galleries without any intervention of the trace maker. Thalassinoides type D is most probably the result of a post-depositional mud infill, in which the dark mud comes from the substrate above. Thalassinoides types C and D are sometimes accompanied by a background of small diameter burrows that form a conspicuous ichnofabric, in which it is impossible to recognize individual ichnogenera. This bioturbated texture shows both a very intense activity of the endobenthic community and a cross-cutting of tiers. Thalassinoides type E. This form corresponds to the ichnospecies Thalassinoides paradoxicus. It is associated only with coarse bioclastic facies, and the form of the burrow shafts and

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Fig. 6. Sketches showing the six different types of Thalassinoides recognized in the studied section. Sample numbers are on the right of the column.

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galleries are controlled largely by the occurrence of shell debris, producing the contorted forms characteristic of this ichnospecies, forming quite an irregular trace pattern. This is indicative of primary substrate control on the assemblage and may be representative of a firm- to hardground environment, in this case the abundance of bioclastic debris. In the same way, the oyster shells of the beds bearing Thalassinoides type E may show borings assignable to Trypanites and Gastrochaenolites (Fig. 6). Thalassinoides type F. This type is associated with eroded bed units, where Thalassinoides galleries are preserved upon the bedding-plane surface. This Thalassinoides type is scarce in the Sal ado section but is very common in other exposures. It developed when the upper part of a substrate bearing Thalassinoides types A, B or C was eroded and the trace level was exhumed. The presence of Thalassinoides type F suggests the development of a local discontinuity or hiatus between the bed with Thalassinoides galleries and the bed deposited above. Arenicolites association This is the rarest association and comprises the ichnogenera Arenicolites, occasionally accompanied by Gordia, Palaeophycus, Planolites, Phycodes, Rhizocorallium and Teichichnus. These traces are largely associated with sandstone beds and even coarser-grained beds, which are uncommon in the Salado section. The Arenicolites association is interpreted as the product of shallower marine conditions and/or a higher level of kinetic energy in the marine substrate (Ekdale et al. 1984). Macrofossil palaeoecology The Vaca Muerta and lower Chachao formations contain a relatively abundant macrofossil fauna, which is nevertheless restricted in diversity (Fig. 7). This comprises nektonic organisms, mostly ammonites, that form the background faunal signal, and a range of benthic faunal associations, discussed below. The palaeoecological information provided was based on the study of bedding-plane associations exposed within the Salado section. Although not strictly quantitative, this was taken to be significant, with due regard taken to the level of autochthony of benthic assemblages. Nektonic organisms The Vaca Muerta and lower part of the Chachao formations in the Rio Salado contain abundant

nekton, mostly in the form of ammonites, although there are rarer accumulations of fish debris. These faunal components are discussed in the following paragraphs. Ammonites. A range of ammonites are known from the Vaca Muerta and lower Chachao formations, and have been illustrated by Riccardi (1983, 1988) amongst others. The majority of ammonites within are preserved flattened, and the original aragonite dissolved. However, some preservation of siphuncles as phosphatic tubes within the shells is observed. Some flattened 'beermat' preservation, with carbonates intact, as well as three-dimensional preservation of ammonites is also recorded from some of the carbonate concretionary horizons at the base of the section. The base of the Salado section (Vaca Muerta Formation) is particularly characterized by ammonites encrusted by a range of byssally attached and cemented bivalves. At the top of the Salado section, within carbonate facies of the Chachao Formation, the 'beermat' preservation is fine enough to preserve ammonite microconchs with intact lappets, and the contents of body chambers, including in situ aptychi, may be observed. This demonstrates little current activity or other shallowmarine processes that would easily destroy or disturb such delicate skeletal features. There are few if any associated benthic bivalves identifiable as either pseudoplankton or benthic colonies associated with these ammonites, and this may indicate an inability of benthic bivalves to colonize foundered ammonites on the seafloor, the most probable conclusion being through the lack of available dissolved oxygen in the bottom waters at this point. The ammonite preservation encountered here has also been identified in the Lower Jurassic Posidonienschiefer by Seilacher et al. (1976), interpreted by him as the product of the early dissolution of the skeletal aragonite, followed by the collapse of the organic periostracum, which forms an organic envelope for the ammonites, and there is no reason to doubt this interpretation here. The presence of epibionts suggests that the ammonites were either the hosts for pseudoplanktonic bivalves (e.g. Seilacher 1960, 1982a, b\ Wignall & Simms 1990; Doyle & Whitham 1991, Etter 1996) or were acting as 'benthic islands' for epizoic colonization (e.g. Cope 1968; Kauffman 1982). Both are recorded from the Salado section. Fish. That there were abundant fish within the Neuquen Basin is indicated through the scattered presence of scales and rare bones, including

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jaws. In many cases these accumulations form cohesive 'lumps' or 'blobs' of bone material that may be safely interpreted as the by-products of predation by other fish, reptiles or other large predators. Elsewhere in the Neuquen Basin the

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Vaca Muerta Formation contains abundant, well-preserved marine reptiles (Gasparini et al. 1997) and, although no large vertebrate material was collected in the present study, this demonstrates the abundance of large nektonic

Fig. 7. Body fossil distribution. 1, Nekton (1A, ammonites; IB, aptycus; 1C, belemnites; ID, fish bones); 2, pseudoplankton (2A, bivalves byssally attached; 2B, oysters); 3, serpulids; 4, gastropods (4A, turretted; 4B, planktonic); 5, byssally attached bivalves (5A, oxitonids; 5B, pectinids; 5C, Entollium; 5D, inoceramids); 6, oysters (6A, Cemtostreon; 6B, Aetostreon; 6C, Deltoidion; 6D, indet.); 7, shallow infaunal bivalves (7A, thin valves; 7B, thick valves); 8, deep infaunal bivalves. A, Nekton; B, pseudoplankton; C, serpulids; D, epibyssate bivalves; E: oysters; F, shallow infaunal bivalves; G, deep infaunal bivalves. Sample numbers are on the right of the column.

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organisms capable of living in a more oxic zone within the water column. Planktonic -pseudoplanktonic associations The planktonic and pseudoplanktonic associations that have been recognized are predominantly molluscan. These are described below. Planktonic gastropod asssociation. Gastropods occur within several levels of the Vaca Muerta and lower Chachao formations. Local monospecific abundances or 'gluts' of small, rounded gastropods were also recorded. These are similar to the gastropod associations recorded by Etter (1995) from the Middle Jurassic of Germany that have been considered to be planktonic (Bandel & Hemleben 1987), but which were thought to be benthic by Etter. In either case, they represent opportunists, either as benthos or nekton, demonstrated by the monospecific nature of the assemblage, controlled as with all monospecific assemblages encounted in the Salado section by primary oxygen levels. Pseudoplanktonic association. This association comprises the close association of epibiontic bivalves with ammonites, the bivalves being found attached to both flanks of the ammonites observed. This is possible because of the 'beermat' preservation, which causes the thicker shelled bivalves to be pushed through the flanks of the ammonite test. The pseudoplankton is restricted mostly to thin-shelled cementing oysters, displaying clear xenomorphism and to byssally attaching, oxytomid-type bivalves. The majority of these bivalves are concentrated within the umbilical regions of the ammonites studied. In most cases the frequency of encrustation is intense, with a majority of ammonites infested, although it should be noted that an accurate count is difficult to determine given there were no bedding-plane surfaces sufficiently exposed to carry out a detailed palaeoecological assessment. Infauna is otherwise restricted, although there are records of some thin-shelled individuals within the section. The pseudoplanktonic association may be distinguished from the otherwise similar benthic island association by: (1) the presence of bivalves on both sides of the ammonites examined; (2) the clustering of bivalves within the umbilical region of the ammonite, and the absence of overgrowth; and (3) the absence of other clusters of bivalves not associated with an ammonite. The pseudoplanktonic association was recorded from several levels within the section studied and, in most cases, no other

associated fauna were recorded from these levels. Where found, associated fauna are generally restricted to isolated bivalves of the same groups as those attached to the ammonites. This represents a classic pseudoplanktonic association given the frequency of the occurence, the presence of epibionts on both surfaces of the ammonites and the overall absence of other benthos. It is similar to those described by Seilacher (1960, 19820, b\ Wignall & Simms (1990), Doyle & Whitham (1991) and Etter (1996), in which settings the presence of pseudoplanktonic host associations in the absence of benthic molluscs is considered typical of low levels of dissolved oxygen at the seabed. Benthic associations A number of benthic molluscan associations may be recognized within the Salado section (Fig. 7). These are recurrent assemblages, and are authochonous or parautochthonous clusters that may be interpreted as approximations of the original skeletal macrofaunal benthos. Benthic island association. The benthic island association is represented by clusters of epifaunal pteriform and ostreid bivalves that are otherwise not abundant. Usually the association is monospecific, comprising clusters of a single epifaunal bivalve taxon typically with an ammonite host. The bivalves are not clustered in specific areas of the ammonite shell, and in the majority of cases the bivalves are seen to be overgrowing the shell margins of the ammonite. Bivalves are only present on the upper surface of the shell. The benthic island association is restricted to relatively few beds within the Salado section, and is easily distinguished from pseudoplanktonic associations that have a much more tightly constrained concentration of bivalves in their umbilical regions. The concept of benthic islands as a threedimensional refuge for benthos, emerging above the lowest levels of dissolved oxygen as developed by Kauffman (1981) for the Posidonienschiefer, is probably too sophisticated for the Salado section, and the level of information available is insufficient to suggest that these islands provided small highs above a stratified lowest oxygen layer. Instead, it is probably sufficient to consider the association as representing marginally increased oxygen conditions, and a relatively soft substrate (see also Wignall 1993). Ammonite inquilinism. An association of ammonites and decapods is recorded on the basis of two specimens of flattened ammonites

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with fragmentary decapod carapaces within their body chambers in the upper part of the Salado section. This is considered to be a primary association for the following reasons: (1) the preservation of the ammonites, which, although flattened, show delicate features such as lappets intact; and (2) the absence of winnowing or other indications of current activity. This suggests a primary, rather than post-mortal, relationship with the decapods finding refuge within the ammonite body chamber on the seabed. The presence of decapods within the body chambers of the sunken ammonites suggests that levels of dissolved oxygen were sufficient to permit life, and this association is in many ways equivalent to that of the benthic island association already discussed. This kind of association has been termed ammonite inquilinism, the post-mortal association of decapod crustaceans within the body chambers of large ammonites. It has been recorded from the Lower Jurassic of England and Germany (Fraaye & Ja'ger 1995). Rotularia association. Serpulids of the genus Rotularia are common at certain levels in the Salado section. These are observed forming small monospecific clusters, although Rotularia is known occurring with other taxa, most notably fragmentary infauna, further up-section. Life orientation of this genus is with the elongate siphonal tube vertical, the main part of the serpulid being oriented within the body of the sediment, and this is the typical orientation of Rotularia within the Salado sediments. Macellari (1984) has discussed the mode of life of Rotularia from the Cretaceous rocks of Antarctica. Three possible modes were suggested: (1) epifaunal suspension-feeders; (2) infaunal deposit-feeders; and (3) infaunal suspension-feeders, with the position of the siphonal tube considered to be of paramount importance in the interpretation of the mode of life. Macellari favoured an infaunal filter-feeding mode with the siphonal tube oriented vertically in life, and this is supported by the present association of rotularids. He also noted the correlation of serpulid abundance with increased mud levels, and recorded the decline in abundance with coarser grained facies. The present association is comparable in that the serpulids found in the Salado section are most commonly located within black muds, with little other associated body fauna. The comparative glut of serpulids associated with certain horizons and in small concentrations is consistent with the perceived characteristics of an opportunistic species (Levinton 1970). In this case, the presence of

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appropriate mud facies is seen to be significant, although it is clear that without suitably oxygenated conditions, at least at dysaerobic levels, it is unlikely that the association would be successful. Entolium association. Shell pavements comprising disarticulated, but otherwise intact, valves of the free-living epifaunal bivalve Entolium are common in the upper part of the Salado section. These epifaunal bivalves are commonly found associated with the tops of carbonate units, and form sparse shell pavements with little or no imbrication and with valves in both stable and unstable orientations. Valves are usually well preserved, but in some beds damage to the valves has occurred, usually in the form of destruction of the wings or shell margin. Entolium is usually interpreted as an freeliving bivalve requiring a firm, or at least not soupy, substrate (Wignall 1990). This interpretation is supported by the relatively high concentration of bioclastic debris within the carbonate units of the Salado section. The concentration of valves most probably represents a shell bed of para-autochthonous bivalves, indicating colonization of a relatively firm substrate with, once again, at least low levels of dissolved oxygen. Other bivalves adapted to softer substrates are typically thin shelled, having adopted the snowshoe approach (Thayer 1975; Wignall 1993; Etter 1996). It is more likely, given the freeliving strategy of these bivalves, that the presence of Entolium is indicative of a firmer substrate. Shallow infaunal suspension-feeding bivalve association. Small, shallow infaunal bivalves are present towards the top of the Salado section, comprising small, thin-shelled and mostly articulated valves of Aphrodina and related taxa. Typically, these taxa have relatively shallow burrows, although they are not associated with rapidly shifting substrates. Thicker shelled Eriphyla, and trigoniids, particularly Steinmanella, are uncommon, but also form a component of the fauna. All of these bivalves are interpreted as shallow infaunal suspension-feeders. The presence of Eriphyla in several of the fine-grained mud-rich or marl units demonstrates the capability of these taxa to inhabit fine-grained, although relatively firm, substrata. The presence of thicker shelled components towards the top of the sequence studied may be readily considered as the product of normally oxygenated bottom waters, and possibly even higher energetic conditions (Wignall 1993; Etter 1996). This

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assemblage is therefore indicative of normal, unrestricted and relatively firm (i.e. non-soupy) substrate conditions. Deep infaunal suspension-feeding bivalve association. Deep-infaunal bivalves are almost wholly associated with the upper part of the section, where they were mostly observed in life position. Commonly, taxa such as Lucinia, Panopea and Thracia occur at the base of carbonate units that often exhibit a bioclastic and strongly reworked top. This is interpreted as the product of strong in situ reworking of more shallow burrowing tiers of the carbonate unit, the level of reworking not tapping into the lower levels inhabited by the deeper burrowing bivalves. Infaunal deposit-feeding association. This is a rare component of the fauna, but, where present, it comprises abundant infaunal depositfeeders composed primarily of small nuculid bivalves. According to Morris (1980), depositfeeders are often associated with dysaerobic conditions, but, as pointed out by Wignall (1993), this may be a function of softer substrates. Exogyrid association. This comprises exogyrid oysters of the genera Ceratostreon and Aetostreon. These associations are common in the upper part of the succession and regularly display a dense concentration of these taxa. These are usually not disarticulated and form beds of up to 1 m in thickness. At their thickest these shell beds are equivalent in facies to those exhibited in the basin central facies of the Chachao Formation, where great thicknesses of alternating Ceratostreon and Aetostreon are found. Ceratostreon is also present as a minor component of the black muds occurring lower down the succession, and, in some cases, they are found with associated serpulids of the genus Rotularia. In this case, small 'clusters' of Ceratostreon are found in an assumed autochthonous association with a 'shell pavement' produced by the concentration of serpulids. This represents a modification of an otherwise soft substrate by the concentration of serpulids, indicating a two-stage colonization process analogous to that demonstrated by Doyle & Whitham (1991) and discussed by Wignall (1993), with similar associations in the Jurassic Oxford Clay of the UK where low-diversity Gryphaea assemblages were considered to be the result of low oxygen conditions, in which an increase in soft substrates combined to reduce the diversity still further. In the case of

the Chachao associations the increased importance of soft substrates is more likely to be a factor where Ceratostreon is found in isolated clusters, rather than in shell beds.

Relative oxygenation curves Following the methodology of Savrda & Bottjer (1986), it is possible to develop a trace fossil model for relative oxygenation of black shales through the relative diversity of trace fossil assemblages. This technique has been reapplied by other authors (e.g. Doyle & Whitham 1991), and has been tested and reviewed by others (e.g. Allison et al. 1995; Etter 1995; Doyle et al 1998; Spalletti et al. 2001), and is generally held to be robust. However, the close association of low oxygen conditions with mud facies has led to difficulties in the recognition of the dominant controlling factor, as outlined by Wignall (1993) and Goldring (1995). Two approaches are used here, that of the standard relative oxygenation curve based on trace fossil assemblages developed by Savrda & Bottjer (1986), and a consideration of relative oxygenation based on body fossil assemblages (Doyle & Whitham 1991; Wignall & Hallam 1991; Wignall 1994). Relative oxygenation from the ichnofauna Following Savrda & Bottjer (1986), two basic characteristics were employed in constructing relative oxygenation curves: (1) the diversity of the assemblage, which for the most part is poorly diverse; and (2) the burrow diameter, including burrow fill. Increasing diversity and burrow diameter equating with increased relative oxygenation. The relative oxygenation curve derived from this approach for the interval studied here is presented in Figure 8. The most significant issue is the transition from laminated to bioturbated shales, with Chondrites being present in each bioturbated interval. Alternation between laminated and Chondritesbearing sediments can be taken as the nominal boundary between anoxic and oxic conditions; effectively the anaerobic-exaerobic/dysaerobic boundary, as defined by Savrda & Bottjer (1986, 1987). Commonly, Chondrites is accompanied by Thalassinoides with small burrow diameters. This is commonest in the black-mud-dominated part of the sequence, the Vaca Muerta and lowermost Chachao formations as developed in the Salado section. Larger burrow-diameter Thalassinoides occur more commonly in the carbonate-dominated parts of the section (Chachao Formation), with a variety

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of fills as described above. This is generally indicative of increasing oxygenation of the basin up-section, consistent with the current models, with a transition from anoxic, through fluctuating

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dysaerobic to oxic conditions determined by the simple relationship of: laminated shalesChondrites/small burrow-diameter Thalassinoides- large burrow-diameter Thalassinoides.

Fig. 8. Oxygen indicators and relative oxygenation curves. A, B and C are suggested oxygenation cycles for the Vaca Muerta Formation. D and E are suggested oxygenation cycles for the Chachao Formation.

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In the lower part of the section, at the transition between the Vaca Muerta and Chachao formations, relatively large burrow-diameter Thalassinoides conforming to types A-D described above occur in carbonate-rich levels between black shales lacking pervasive bioturbation, with the whole part of the section lacking infaunal bivalve associations. This represents a paradox, for, as explained above, the presence of Thalassinoides has been benchmarked in most palaeo-oxygenation studies as an indicator of oxic conditions, at odds with the absence of other benthos and a background of pseudoplanktonic bivalves attached to nektonic ammonites. Relative oxygenation from body fauna Benthic and pseudoplanktonic fauna are a valuable proxy for the development of relative oxygenation, although not as widely employed as the trace fossil models. In general, the relationship of abundance and diversity is seen as critical, as with most assemblages it is the transition from opportunistic to equilibrium species (Levinton 1970; Morris 1980; Doyle & Whitham 1991; Wignall & Hallam 1991). The presence of benthos with a pseudoplanktonic habit is also significant (Wignall & Simms 1987; Doyle & Whitham 1991). In the Salado section the presence of nekton ammonites - is pretty well universal, with ammonites reasonably abundant throughout the section, declining to the top of the section. Pseudoplanktonic bivalves are similarly distributed, but with a greater abundance in the mid-part of the section. It is the relationship between the epifaunal molluscan fauna - epibyssate bivalves and oysters - and the infaunal components that is seen to be significant in terms of relative oxygenation. Epifaunal assemblages include the benthic island associations, and the concentrations of gastropods and Entolium. Similarly, isolated clusters of exogyrid oysters are significant, and may represent increasing oxygen (dysaerobic conditions) and fluctuation in substrate from soft to firmer. The serpulid Rotularia is taken to represent an opportunistic species and is common at the base of the succession, occurring beneath the first common pseudoplanktonic development on the nektonic elements. It may well also carry a primary low oxygen signal, but it is thought to prefer softer substrates. The vast majority of the benthic molluscs present are suspension-feeders, with relatively few deposit-feeding bivalves present. Like that

of the serpulid Rotularia, the palaeoenvironmental signal from these deposit-feeders is ambiguous, with low oxygen being overprinted by a preference for soft substrates. Epifaunal bivalves unattached to nekton occur only at the base and very top of the succession, and are generally exclusive of the infaunal bivalves. Shallow infaunal bivalves are seen within all the major facies, and in some cases occur exclusive of epifaunal bivalves. The presence of deep infaunal bivalves is taken to represent the maximum oxygenation state present, and these molluscs are common only in the carbonate-rich basal Chachao Formation (Fig. 8). As described by Wignall (1993) increasing diversity and burrow depth can be a function of both increasing oxygenation, and increasing substrate suitability, and therefore difficult to disintangle. Despite this, the overall signal is in line and in keeping with what could be expected from the lithofacies and the ichnological information - increasing oxygenation up-section, associated with the increase in carbonate facies (Fig. 8). Discussion In the main, the relative oxygenation curves produced by both datasets agree well, particularly in the mid-upper part of the succession, characterized by an increase in carbonate units. The basic trend in this upper part of the succession is a cyclical arrangement with increasing oxygenation of bottom waters indicated by the cycle: laminated Chondrites- small-diameter Thalassinoides—large-diameter Thalassinoides Arenicolites-Thalassinoides with bioclastic infill. This ichnological cycle is matched with a similar cycle of body fossil occurrence: pseudoplankton-epifaunal bivalves-shallow infaunal bivalves. This is, in itself, relatively unremarkable, demonstrating a cycle of increasing oxygenation of bottom waters, allowing for a more prolonged colonization window. The cycles themselves overprint the pattern of marl-packe/wackestones observed, with no significant relationship observed between increased faunal diversity and increasing bioclastic carbonate. However, on the broad scale, the observed cycles are mostly associated with the upper carbonate-rich facies, overlying the otherwise low-oxygen black-shale-dominated facies at the base of the section. At least three of these cycles can be recognized in the carbonate-rich section (C-E in Fig. 8). The basal black shales do, however, show some exception to the general pattern, and, in some cases, there appears to be conflict

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between trace and body fossil oxygenation signals. Within the black shales there is at least one, and possibly two, further laminated- Chondrites/ small Thalassinoides large Thalassinoides cycles (A and B in Fig. 8) that may possibly be substrate controlled, in that the large Thalassinoides assemblages are seen associated with thicker carbonate layers. The basal cycle B is not in any way mirrored by the body fauna, which lacks either epi- or infaunal bivalves (Fig. 8). The cycle below (A) is associated with common serpulids and, interestingly, the colonization of the carbonate components with shallow, suspension-feeding, infaunal bivalves. This is consistent with welloxygenated bottom waters, and is inconsistent with the relatively low diversity of the trace fauna. As previously recognized by Wignall (1993), and underlined by this study, disentangling the substrate and limited oxygentation signals in black shale facies can be a difficult task, especially where there is a definite preference for softer substrates (e.g. Rotularia, Chondrites, deposit-feeding bivalves) that could be misinterpreted as a low oxygen signal, where diversities are low. As such, this study underlines the importance of the comparison of more than one dataset in the interpretation of relative oxygenation. Clearly, the discordance between body fossil and ichnological factors demonstrates that substrate can be a limiting factor as important as that of dissolved oxygen levels.

Conclusions The Tithonian-Valanginian succession exposed in the Salado valley shows the transition from black shale to carbonate facies, although the later biohermal carbonates are limited and largely undeveloped. Our analysis of this section illustrates the following important points: (1) that carbonate production increases up-section; (2) that faunal associations are mostly limited to poorly diverse epibenthos and pseudoplankton in the lower part, with increased diversity in the lower Chachao section, including shallow and deeper infaunal bivalves; (3) that there is a background level of laminated shales to Chondrites bioturbation typical of anoxic-suboxic conditions; (4) that Thalassinoides suevicus is commonly developed in micritic limestone and carbonated sands units throughout the section. Interpretation of the palaeo-oxygenation of this section on the basis of trace fossil taxa alone (cf. Savrda & Bottjer 1986) would give an inaccurate interpretation because of: (1) widespread Thalassinoides burrows of more

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than 1 cm in diameter at the base of the section suggesting at least upper dysaerobic facies, whilst the coeval, greatly reduced, and probably opportunistic fauna of rotularids and pseudoplanktonic bivalves on ammonite hosts is indicative of lower dysaerobic or exaerobic facies; (2) the almost exclusive development of Thalassinoides in limestones and sandstones, while associated body fauna are not as facies dependent. We believe that there was a primary substrate control on trace fossil diversity and occurrence, with a primary oxygenation signal provided by body fossil evidence. This study demonstrates the importance of integrated trace and body fossil analysis in the fuller understanding of black shales. When comparing the curves inferred from fossils and trace fossils in the Vaca Muerta and Chachao formations the mismatch between the palaeo-oxygenation curves inferred sedimentologically and those obtained exclusively from fossil traces (Savrda & Bottjer 1986, 1987, 1989) indicates that trac fossils should not be used in isolation. This work was funded by a CONICET-Royal Society exchange grant, the CONICET Research Project PID 858/98 and SETCYP Project PICT 07-08451, which we gratefully acknowledged.

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Geochemistry and petrology of a Middle Tithonian limestone-marl rhythmite in the Neuquen Basin, Argentina: depositional and burial history R. A. SCASSO1, M. S. ALONSO1, S. LANES1, H. J. VILLAR1 & G. LAFFITTE2 1

Departamento de Ciencias Geologicas, Facultad de Ciencias Exactas y Naturales,

Universidad de Buenos Aires, Ciudad Universitaria, Pabellon 2, 1° Piso, 1428 Ciudad de Buenos Aires, Argentina (e-mail: 2

[email protected])

M&P Systems, Mir6 525, 1428 Ciudad de Buenos Aires, Argentina

Abstract: The Middle Tithonian Los Catutos Member (Vaca Muerta Formation, Neuquen Basin), is lithologically and geochemically similar to limestone-marl alternations from the Late Jurassic of the northern hemisphere. Both marls and limestones are pelbiomicrites and biopelmicrites principally composed of pellets, radiolaria, forams, ostracods, equinoids, spicules of sponges and gastropods, cemented by several generations of calcite cement. Smectite and interlayers are the major epiclastic components of the fraction below 2 (xm, reflecting pedogenic processes developed on volcanogenic source rocks. More abundant kaolinite in some marls reflects stronger humid conditions in the source area and enhanced terrigenous supply. A12O3 content is demonstrated to be a reliable indicator of clastic input. The same is not true for silica, often related to high biogenic productivity of siliceous organisms. Rocks show total organic carbon (TOC) contents up to 1.95% and constitute regular to good sources for hydrocarbons, although thermally immature. Rhythmites formed gently sloping mounds accumulated in a distal submarine ramp under low-energy and poorly oxygenated open-sea conditions. Sedimentation rates were high due to high productivity on the sea surface, and supply of terrigenous and carbonate sediments transported by suspension plumes originated in shallow, photic areas. 613C values correspond well with the published curves for the Tithonian sea water and with other records from Tethyan limestones. A preliminary analysis of negative excursions of 513C point to a productivity crisis or a mixture of water layers in a stratified sea with a periodicity of 400 ka, which could be a result of changes in the orbital eccentricity of the Earth. Light isotopic composition of O in bulk rocks is the result of diagenetic neomorphism and cement precipitation. Calculated palaeotemperatures from 518C are consistent with those derived from measured vitrinite reflectance (Ro%) and burial history reconstruction. Data indicate initial burial during the Tithonian extending up to the Lower Cretaceous, a short period of uplift (Intravalaginian tectonic phase), and renewed uplift during the Cenomanian followed by significant Late Cretaceous sedimentation and Pliocene thrusting.

Rhythmic successions of limestones and marls of different ages and sedimentary environments are frequent in the geological record (Einsele & Ricken 1991; Hemleben & Swinburne 1991). In spite of their monotonous appearance, these rocks show particular geochemical signatures and microscopic features that result from the varying sources of their biogenic and non-biogenic components and from their burial-diagenetic history. A particular type of marl and limestone rhythmite, known as Tlattenkalke' or lithographic limestone (Barthel et al. 1994), is composed of tabular beds, several centimetres thick, and bounded by very regular and equally spaced bedding planes. Some lithographic limestones

(i.e. Solnhofen, Germany) are well known due to the excellent preservation of their fossils that resulted from rapid burial in a low-energy, hypersaline and/or anoxic shallow carbonate platform. These beds are known as 'fossil lagerstatten' and they can even preserve the soft parts of the organisms. The worldwide distribution of lithographic limestones in the Upper Jurassic (Dehm 1956) is remarkable. Cyclically alternating light brown marls and limestones from the Tithonian Los Catutos Member of the Vaca Muerta Formation (Leanza & Zeiss 1990) crop out near Zapala, in the southern part of the Neuquen Basin (Fig. 1). They are massive and contain well-preserved

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 207-229. 0305-8719/05/$15.00 © The Geological Society of London 2005.

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Fig. 1. (a) Location map of the Neuquen Basin, (b) Late Middle Tithonian palaeogeography (from Legarreta & Uliana 1996; Cruz et al. 1999). Local source areas are thought to be present in the Dorsal Neuquina area.

fossils (Leanza & Zeiss 1992). According to their tabular geometry, thin bedding and fossil preservation, some limestones from Los Catutos Member can be considered as 'lithographic' (Leanza & Zeiss 1992; Scasso et al. 2002). The purpose of this paper is to analyse the depositional and burial history of the Los Catutos Member on the basis of its sedimentary petrology, organic and inorganic geochemistry (major and trace elements, and stable isotopes) and stratigraphic framework. In addition, this is the first example of a southern hemisphere limestone-marl rhythmite studied in this way. Comparison with sequences of the northern hemisphere will allow global palaeoclimatic reconstructions for the Late Jurassic. Methods Five sedimentological sections were described and sampled bed by bed in the 25 m-thick

rhythmite in Los Catutos, and other quarries of the Loma Negra Company in the surrounding area. Hard limestone and recessive marls were characterized by their resistance to erosion in the field. The field logs were then correlated with the chemical composition of each limestone and marl in the sequence (see Bausch 1997). All the sections are located in a small area of about 4 km2. Thin sections of limestones and marls were studied under a petrographic microscope. Components were semi-quantitatively estimated by comparison to graphic charts (Schafer 1969; Dietrich et al. 1982, data sheets 15.1 and 15.2). The main mineral phases in the limestones and marls were determined by X-ray diffraction (XRD) on bulk, grinded samples. Preliminary studies on siliceous microfossils were performed on the insoluble residue (Kiessling pers. comm. 2004). Chemical analyses of samples were performed by alkaline fusion with Na2CO3 followed by acid

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dissolution of the residue; K2O and Na2O were analysed by FA AS (flame atomic spectroscopy), MnO and MgO by AAS (atomic absorption spectroscopy), A12O3 and Sr by ICP-OES (inductively coupled plasma-optical emission spectrometry), TiO2 and P2O5 by AS (absorption spectrometry), CaO by gravimetric and volumetric method and SiO2 by a combined gravimetric-AS (absorption spectrometry) method. In addition, loss of ignition (LOI), loss of humidity at 110 °C and a gravimetric method for CO2 were performed on separate samples. Clay minerals were studied in 37 samples. Siliciclastic components were concentrated by dissolving limestones and marls in a weak acid (monochloracetic acid). Terrigenous content was determined by weighting the insoluble residue and the <2 jjim fraction was separated using an Atterberg tube. XRD analyses were run on untreated, glycolated and heated (550 °C) oriented samples. A Philips 1130 diffractometer was used under the following operating conditions: Cu/Ni, radiation, range: 1 x 103 counts per second and scanning speed 2° 20 min"1. Percentage evaluations were based on peak areas, corrected by empirical factors on glycolated samples. 3.57/3.54A lines were used to differentiate kaolinite and chlorite reflections. Identification of mixed layer components was made after Moore & Reynolds (1989). Bulk-rock isotopic composition of limestone and marls were measured in two sections (ZAP B and ZAP E) with a mass spectrometer Finnigan Mat 252 on-line with a Carbo-Kiel device for sample preparation in the University of Erlangen

(reaction temperature 75 °C). Average accuracy precision was ±0.05%0 for 513C and 0.08 for 5180. Total organic carbon (TOC%) was determined with a Leco EC-12 carbon analyser after treating the samples with concentrated HC1 to remove carbonates. Kerogen type was determined by means of a Rock-Eval II pyrolysis instrument, which also provided the rmax parameter for thermal maturity estimation. Visual kerogen analysis was performed both on 'palynological' slides and on polished epoxy plugs of kerogen concentrate, using combination of white light and UV/blue light ('fluorescence' mode). Thermal maturity assessment was carried out by measuring vitrinite reflectance (Ro%) following standard methods. In addition, values obtained were cross-checked with TAI (thermal alteration index) estimations and fluorescence signatures of the kerogen. Burial and thermal history was reconstructed using a one-dimensional (ID) basin modelling program, the Genex (Beicip-Franlab) version 3.4.0. Geological data input (Table 1) was derived from available geological maps and published subsurface information of neighbour areas (Ploszkiewicz et al. 1984; Mitchum & Uliana 1985; Leanza & Zeiss 1990; Legarreta & Uliana 1996, 1999; Leanza & Hugo 1997; Cruz et al. 1999, 2000; Zavala 2000).

Geological setting The Neuquen Basin was a wide SE-NW-trending gulf (Engolfamiento Neuquen-Aconcagua,

Table 1. Database for burial history and thermal maturation calculation. Ages for the beginning and end of the accumulation, as well as thickness, are from local and regional studies in the right-hand column

Age

Beginning (Ma)

End (Ma)

early late Tithonian

135

133

94

133

130

225

Bajada ColoradaMulichinco

late Tithonian Berriasian BerriasianValanginian

130

121

450

Gently uplift Agrio-Centenario La Amarga

early Hauterivian HauterivianBarremian

121 118

118 113

-100

Intense uplift

late Barremianmid Albian CenomanianConiancian Santonian - Recent

113

95

-1200

95

80

700

80

0

-769

Formation Upper Vaca Muerta Picun Leufu

Rayoso— G.Neuquen Allen, Jagiiel, Roca and younger beds

Thickness (m)

600

Source Leanza & Zeiss (1990) Leanza & Zeiss (1990); Leanza & Hugo (1997) Leanza & Zeiss (1990); Leanza & Hugo (1997); Zavala (2000) Legarreta & Uliana (1999) Cruz et al. (2000); Legarreta & Uliana (1999); Leanza & Hugo (1997) Legarreta & Uliana (1999); Cruz et al. (2000) Legarreta & Uliana (1999); Cruz et al. (2000) Leanza & Hugo (1997)

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R. A. SCASSO£rAL.

Legarreta & Uliana 1996) connected to the palaeo-Pacific Ocean at its NW end (Fig. 1) and filled with a thick Late Triassic- Palaeogene sedimentary succession. During the Late Jurassic the Neuquen Basin was located between the South American subsiding foreland to the east and a volcanic arc to the west (Legarreta & Uliana 1996). Basin stratigraphy of the Argentine part of the Neuquen Basin was reported by Digregorio & Uliana (1980), Leanza (1981) and Gulisano et al (1984), whereas Hallam et al (1986) compared the stratigraphy of the Argentine and Chilean parts of the basin. Sedimentation began with a rift phase during the Triassic, which evolved to an early sag stage by the Late Jurassic-Early Cretaceous (Mitchum & Uliana 1985), when the Vaca Muerta Formation was deposited. The general sequence stratigraphic framework of the Neuquen Basin was published by Legarreta & Gulisano (1989), while Mitchum & Uliana (1985) and Legarreta & Uliana (1996) focused on the Upper Jurassic-Lower Cretaceous sequence stratigraphy, correlating the transgressions and regressions recorded in the basin with the global eustatic changes. Spalletti et al. (2000) subdivided the Tithonian from the southern Neuquen Basin into three sequences mainly composed by transgressive and highstand systems tracts deposited in a mixed carbonatesiliciclastic ramp. During the Jurassic and Early Cretaceous, the Neuquen Basin was flooded by an epeiric sea connected to the Pacific Ocean. The sea reached its maximum level in the Tithonian Early Berriasian, when alternating black shales and limestones of the Vaca Muerta Formation were deposited. These fine-grained facies interfinger with near-shore coarse elastics and carbonates towards the basin margins (Leanza & Hugo 1977; Legarreta & Gulisano 1989; Spalletti et a 2000). In the Andean Range, Tithonian carbonate-bioclastic rocks correlate with conglomerates, ignimbrites and lava flows (Hallam et al. 1986; Cegarra et al. 1993), and were deposited in narrow carbonate platforms around eruptive centres (Sanguinetti 1989). In the southern part of the Neuquen Basin, carbonate sedimentation was mostly hindered by an abundant siliciclastic influx from the SE. However, two carbonate units, the Los Catutos Member of the Vaca Muerta Formation and the Picun Leufu Formation, were accumulated in the Zapala-Los Catutos area as a part of a 420 m-thick Tithonian succession. The Piciin Leufu Formation concordantly overlies the Vaca Muerta Formation and records the transition to the Berriasian (Leanza & Hugo 1977; Leanza & Zeiss 1990).

The study area is located on the SE part of the Sierra de la Vaca Muerta Anticline (Lambert 1956), close to the village of Los Catutos. Locally, the limestone and marl beds form a wide syncline that passes towards the NE into a narrow anticline whose eastern limb hosts the El Ministerio Quarry, the type locality of Los Catutos Member. There, a 70 m-thick succession of alternating limestones, marls and mudstones (Fig. 2) is exposed (Leanza & Zeiss 1990). The Los Catutos Member at El Ministerio Quarry includes a fossil assemblage of cephalopods, marine reptiles and fishes (Cione et al. 1987; Gasparini et al. 1987) and five ammonit levels (Leanza & Zeiss 1990, 1992). The lower three ammonite levels correlate well with the Windhauseniceras internispinosum Zone (Uppermost Middle Tithonian), although the lowest one might be slightly older, reaching the Aulacosphinctes proximus Zone (mid Middle Tithonian, Fig. 2). The uppermost ammonite level reaches the Middle Tithonian-Late Tithonian boundary. The part of the section composed of cyclically arranged limestones and marls coincides with the Catutosphinctes rafaeli Subzone (Leanza & Zeiss 1992) of the lower part of the uppermost Middle Tithonian. Nannofossils match the NJ20b Zone confirming a Late Middle Tithonian age (Scasso & Concheyro 1999).

Los Catutos section The 25 m-thick, rhythmically alternating limestones and marls (Figs 2-4) are well exposed at the Los Catutos and Loma Negra quarries. The limestone-marl rhythmites overlay a 164mthick siliciclastic succession of dark-brown and grey coloured mudstones with sporadic thin intercalations of yellowish-brown and grey marls containing fragmented ammonites. The uppermost 30 m of the succession are well exposed in the Los Catutos Quarry. The rhythmite interval is, in turn, covered by 15 m of lenticular mudstones and marl beds, filling a well-exposed palaeochannel in the central quarry of the Loma Negra Company. The top of the Los Catutos Member is marked by two isolated limestone beds (Fig. 2) cropping out between recessive unconsolidated sediments (probably mudstones). Limestones are massive to faintly or well laminated and have a yellowish-brown colour on weathered surfaces, and are greenish or bluishgrey on fresh surfaces. Beds average 0.3-0.4 m in thickness, although they can exceed 1 m towards the top of the succession. Marls are usually faintly laminated and thicker near the base. They average 0.15-0.20 m (Fig. 3), and reach a maximum thickness of 0.5 m. Thin

Fig. 2. Schematic section of the 25 m-thick limestone and marl sequence sampled in the sections ZAP B and ZAP E+D of Los Catutos area. Nannoplankton zone NJ20b (Scasso & Concheyro 1999), ammonite levels (X+a, X, Y, Z, W) and ammonite zonation (Leanza & Zeiss 1992) are also indicated.

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Fig. 3. View of the lower part of Los Catutos Member, showing recessive marl beds up to 50 cm thick (average 15-20 cm) interbedded with hard limestone beds averaging 30-40 cm in thickness. Loma Negra Norte quarry.

marl beds, 5 cm thick or less, appear interbedded with thick limestone beds at the top of the succession (Fig. 4). Collapsed and technically deformed ammonites are abundant in the lower part and at the top of the rhythmite interval. The phragmacones lay parallel to the bedding planes, and show the last whorls and chamber collapsed along fractures running parallel to the outer limit of the shell (Fig. 5). In addition, they are tectonically stretched towards 65°. Diffusive bioturbation similar to Chondrites isp. or irregular mottling was rarely distinguished in otherwise massive limestone and marl beds. Laminated limestones and marls show alternating dark and light laminae separated by sharp irregular-slightly wavy bedding planes (Fig. 6a, b). Laminae (1-3 mm thick) thin out laterally resembling a flaser lamination. Delicate lamination is often emphasized by the orientation of particles parallel to the laminae. Light laminae reach a maximum thickness of 10 mm and often show normal grading. Dark laminae contain abundant micrite or pellets and some subtle seams of iron oxides; they can attain a maximum thickness of 5 mm, mantling the irregularities of the light-coloured laminae. Petrography Limestones and marls are calcareous mudstones (Dunham 1962). X-ray diffraction only

demonstrates the presence of calcite, almost the exclusive component in the carbonate fraction, and very scarce dolomite. According to their microscopic modal composition and internal fabric (Folk 1959) four lithological types were recognized: (1) laminated biopelmicrites (Fig. 6a, b); (2) massive biopelmicrites; (3) laminated pelbiomicrites; and (4) massive pelbiomicrites. Pelbiomicrites and biopelmicrites contain skeletal grains in a pelletoidal or micritic matrix, rarely cut by veins of blocky spar calcite (Fig. 6b, c). The pelletoidal matrix is composed of micritic elliptical particles (most probably faecal pellets) showing tangential and linear interpelletoidal contacts, and pore spaces filled with clean xenotopic microspar and spar. Bioclasts are dominated by planktonic, with minor nektonic specimens (radiolarians, planktonic foraminifera and scarce ammonites) together with less abundant benthic foraminifera (similar to Lenticulina sp. and Quinqueloculina sp.), ostracods, echinoids, sponges, bivalves and gastropods. Pellets, peloids, phosphatic particles and terrigenous grains are the other components. Rhaxes (kidney-shaped siliceous sponge spicules with small mamellons on the surface that cannot be distinguished from spumellarians in thin section) and radiolarians (mostly Parvicingula-type acuminate nasselarians, as

TITHONIAN LIMESTONE-MARL RHYTHMITE

213

Fig. 4. View of the upper part of Los Catutos Member in the central Loma Negra Quarry, showing up to 1 m-thick limestone beds interbedded with thin marls. The uppermost limestone bed is covered by green-brown clastic mudstones of the Upper Vaca Muerta Member. To the right a large cut-and-fill sedimentary structure is eroded in the mudstones and filled with fine-grained, gently dipping mudstones and marls. The size of the sets and the angle of dipping decrease to the right as a result of the gradual filling of the channel structure.

shown in Fig. 6b) are abundant (4-20%, diameter <0.57 mm) and more common in the massive than in the laminated limestones. They are flattened or fragmented, partially to wholly replaced

Fig. 5. Deformed ammonite mould with the phragmacone laying parallel to the bedding plane. Note that the last whorls and chamber collapsed along fractures (arrowed) running parallel to the outer limit of the shell. The mould is stretched towards the Az 65° (long arrow) indicating a maximum compression from Az 155° (NNW-SSE) in Los Catutos Quarry.

by calcite and rarely preserving their original siliceous composition (Fig. 6c, d). They are usually rimmed by micritic envelopes, although in the massive biopelmicrites they can show syntaxial overgrowths, irregular microspar rim cements or a first generation of equant spar rim cement coated by a micritic envelope. Ostracods (2-4%) equal or less than 0.2 mm in diameter are disarticulated, fragmented and oriented parallel to the lamination. Only one ostracod was found articulated, filled with drusy spar, with both valves calcitized. In the laminated pelbiomicrites they are often replaced by blocky spar or micritized. Foraminifera (1-6%, 0.25-0.5 mm long) belong to planktonic (Figs 6c, f & 7g) and benthic groups, the latter represented by uniserial, biserial, planospiral coiled and milioliid coiled specimens. They mainly appear micritized, although in the massive biopelmicrites they can be calcitized with the chambers always filled with drusy spar. Echinoid remains (2-20%, diameter about 0.5 mm) include spines (Fig. 7e, f) and plate fragments (Fig. 7a-c) usually rimmed by micritic envelopes. The spines can be surrounded by an irregular and discontinuous sparitic cement fringe. In the laminated biopelmicrites the

214

R. A. SCASSOErAL.

Fig. 6. (a) & (b) Close views of a laminated pelbiomicrite (lamination polarity towards the top of the photography). Particles parallel to the laminae, irregular base and normal grading of the light-coloured laminae can be seen. Scale bar is 10 mm; x 2.5, plane-polarized. In (a) the arrow points to the echinoderm plate in Figure la. In (b) a vein of blocky spar calcite and a peloid (arrow) can be seen, (c) Radiolarians replaced by calcite (bigger, white spots). Arrows point to pellets that are often separated by blocky sparite filling interpelletoidal pores. Scale bar is 1 mm; x 10, plane-polarized, (d) Detail of the radiolarians in (c). The smaller one to the right is partly replaced by blocky calcite, the other one to the left is completely replaced. Scale bar is 0.5 mm; x 20, cross-polarized, (e) Planktonic foraminifer surrounded by a micritic envelope. Arrow points to a calcitized pelecipod valve. Scale bar is 1 mm; x 10, plane-polarized, (f) Same foraminifer as in (f) showing chambers filled by spar. Scale bar is 1 mm; x 10, cross-polarized.

echinoderm fragments show syntaxial overgrowths coated by micritic envelopes, while in the laminated pelbiomicrites the inverse order of cements is found. Sponge spicules mostly replaced by calcite (1-8% up to 0.5 mm long) only appear in the laminated biopelmicrites and pelbiomicrites being monoaxonic, biaxonic, triaxonic and hexaxonic.

Bivalves (1-12%, diameter <3.7 mm) appear in the laminated biopelmicrites, laminated pelbiomicrites and massive biopelmicrites. They are disarticulated, fragmented, slightly to well rounded, sometimes ornamented by strong ribs (Fig. 7d). Shells usually underwent neomorphic replacement by drusy spar, although some calcitized specimens have preserved a relictic fibrous or columnar microstructure. In one sample of

TITHONIAN LIMESTONE-MARL RHYTHMITE

215

Fig. 7. (a) Echinoderm plate surrounded by micritic envelope and radiolarian (r), x 10, plane-polarized, (b) Fragmented echinoderm plates, x 10, plane-polarized, (c) Detail of a fragmented echinoderm plate showing a syntaxial overgrowth (arrow), x20, plane-polarized, (d) Calcitized pelecipod valve, x5, plane-polarized, (e) & (f) Longitudinal section of an echinoderm spine showing a partial micritic envelope, x 20, plane-polarized in (e) and crossed polarized in (f). (g) Blocky spar patchly replacing the matrix of a laminated biopelmicrite. Planktonic foraminifer (arrowed) surrounded by a thick micritic envelope, x 10, plane-polarized, (h) Similar to that in (g) showing neomorphic sparite mosaic (base 'layer') and non-replaced pellets (top layer), x20, plane-polarized. Scale bars are 1 mm in all the photographs.

laminated biopelmicrites, a fragment from the ligament area of an inoceramid was identified. Probably most of the prismatic fragments of the bioclastic fraction are derived from valve breakage. Ammonites (1-3% 1.5-2 mm long) occur in the laminated biopelmicrites and pelbiomicrites. They are fragmented, wholly replaced by calcite with the chambers filled with drusy spar.

Gastropods (1%, 1.5-2 mm long) only occur in the laminated biopelmicrites as fragmented and wholly neomorphized specimens with their chambers filled with drusy spar. Phosphatic particles (1-10%, diameter < 0.4 mm) are acicular or spheroidal in shape showing a massive or fibrous texture. Among the non-skeletal particles, pellets, peloids and terrigenous grains are identified.

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R. ASCASSOETAL

Pellets (15-85%, diameter of 0.09-0.19 mm) occur in all lithological types, being less abundant in the laminated biopelmicrites and pelbiomicrites. They are micritic, elliptical or spheroidal, often clumped in aggregates and most probably of faecal origin. Scarce, irregular-shaped and rounded, micritized peloids (0-5%, maximum diameter about 0.5 mm) appear in the biopelmicrites and laminated pelbiomicrites. They are of larger size than the pellets and their origin is uncertain. Terrigenous grains (0-4%) comprise feldspar, pyrite, volcanic fragments, biotite, muscovite, chlorite, glauconite and mudstone intraclasts. Feldspars include subrounded plagioclase and orthoclase partly replaced by blocky calcite and/or partly altered into kaolinite. Minor framboidal pyrite is partially altered into iron oxides. Lithic fragments of andesite show pilotaxic and felsitic textures. Unaltered flakes of muscovite, biotite and chlorite, together with rare glauconite grains and rounded intraclasts of bioclastic mudstones, are minor components of the rocks. The greatest proportion of terrigenous, pellets, echinoderms and phosphatic particles occur in a sample of laminated biopelmicrite. The maximum percentage of bivalves also appears in the laminated and massive biopelmicrites, and in one of the laminated pelbiomicrites. Sponge spicules and phosphatic particles are most common in the massive biopelmicrites. Five types of cements were identified. In the laminated pelbiomicrites, massive pelbiomicrites and laminated biopelmicrites the cement succession is as follows: (1) continuous or discontinuous micritic envelopes, preferably around bioclasts (Figs 6c, e & 7g); (2) drusy spar fillings, calcitization and replacement of radiolarians and bivalves (Fig. 6c, d); (3) blocky spar calcite filling intergranular spaces or patchy replacing the pelletoidal or micritic matrix (Fig. 7g, h); (4) rare, euhedral dolomite developed only in the matrix; and (5) irregular aggregates of iron oxides filling seams or fine fractures. Most of the bioclasts show dissolution of hard parts before replacement by drusy spar (neomorphic replacement), whereas some others show preservation of the relictic valve structure. Drusy fillings of empty chambers are usual. Dolomite content is usually constrained to a few, 0.01 mm-long, non-mimic (sensu Sibley and Gregg 1987) and non-pseudomorphic crystals replacing the micritic matrix. Another cement succession is recognized in the massive biopelmicrites. The earliest cement generation includes discontinuous rim cements of blocky calcite spar exclusively around radiolarian, together with scarce syntaxial overgrowths in echinoderms and radiolarians (Fig. 7c). The

second cement generation comprises continuous or discontinuous micritic envelopes. The third cement generation is made up of small drusy spar filling interpelletoidal spaces (Fig. 6c), while the fourth comprises xenotopic drusy spar (neomorphic calcite?) that concentrates towards the bed tops. Finally, the fifth cement generation includes iron oxides in seams. Chemical composition of limestones and marls The section in the central quarry of Loma Negra Company (ZAPB, Fig. 4) shows CaCO3 content between 74 and 82% in the limestones, and between 70 and 78% in the marls. At ZAP B and ZAP E-D sections (Figs 3 & 4) the amount of SiO2 varies between 10 and 19% in marls and between 10 and 16% in limestones; the A12O3 percentage ranges from 2 to 5.9% in marls, and from 1 to 5.6% in limestones (Fig. 8). The Sr contents are relatively high (800-llOOppm) but not rare for Jurassic limestones with abundant bioclasts. Recessive marls are always richer in A12O3 than the immediately underlying and overlying hard limestone beds. SiO2, A12O3 and CaCO3 amounts of both lithologies overlaps if the entire section is analysed, because of the low average CaCO3 content in both limestones and marls, from the lower part of the section. This is not unusual because the carbonate content of the limestone-marl boundary ranges from 65 to 85% (Seibold 1952). Clay minerals Expandable minerals dominate throughout the unit, be they rather pure smectite or mixed layers illite-smectite with variable amounts of expandable layers. Kaolinite is present frequently as a trace, but in some specific levels is the dominant component. Illite and chlorite only appear as discrete mineral phases in few samples. Except for the noticeable presence of kaolinite in some marl beds, no clear patterns of distribution of minerals are observed between the limestones and marls. Isotopic composition of limestones and marls 518O range between —3.5 and — 5.5%c for limestones and marls (limestones average -4.87%o and marls — 4.43%c), and 613C ranges between — 1 and +1.55%c (limestones average 0.93%c and marls 1.025%o). 513C is about + 1.40%o, punctuated by negative excursions along the section (Fig. 9), whereas 518O tends to decrease towards

TITHONIAN LIMESTONE-MARL RHYTHMITE

217

Fig. 8. ZAP D section showing good correlation between lithology (as identified in the field) and alumina content. L, limestones; ML, marly limestones; LM, limy marls; M, marls.

the top. Distribution of both parameters is roughly covariant (Fig. 10), although covariance is especially clear in the upper part of the column. Organic matter Organic content and Rock-Eval pyrolysis. The organic content is low-fair to moderately high, with TOC values spanning the range 0.201.95%. Rock-Eval pyrolysis data point to variable qualities of the organic matter, basically indicating II-II/III kerogen types (Fig. 11). Hydrogen indices are around 450-650 mg g"1, but are accompanied by widely variable oxygen indices (range 40-180mg g"1) that suggest a primary variation of the kerogen type or its preservation during deposition or, alternatively, a kerogen alteration process in the outcrop (weathering). This last process is suggested by the comparative TOC and Rock-Eval analysis in two different portions of the same sample (ZAP A-22-a, 'altered' yellow, outer rim, v. ZAP A-22-b, 'preserved', grey, inner part). From the viewpoint of the hydrocarbon habitat, the overall data of the organic-rich samples suggest fair-good source rocks for oil and gas generation. However, general RockEval and, in particular, the rmax data (see below) are consistent with a low thermal maturity and, therefore, indicate low convertibility. Kerogen microscopy. The visual kerogen analysis shows strong predominance (90-95%) of unstructured material displaying heterogeneous, mostly gold-orange, medium-intense fluorescence; highly-fluorescent, yellow sectors

are also recognized. Structured particles are scarce, mostly representing small sphere bodies (possibly algae); identifiable terrestrial relics predominantly constituted by fusinite and semifusinite particles also occur. Some solid bitumen participation was also observed. The organic association is comparable to that documented for the Vaca Muerta Formation in neighbouring areas of the basin (Cruz et al 1999), dominantly derived from marine, likely algalbacterial organic matter, which constitutes the source of the liquid hydrocarbons of numerous oil fields of the area. The 'preserved' sample contains approximately 25-30% framboidal to finely disseminated pyrite, whereas the 'altered' sample contains no more than 5-10%. Thermal maturity. The Rock-Eval data (rmax in the range 422-427°C), along with the microscopy data (Ro, 0.41-0.48%; moderateintense fluorescence), suggest that the samples are immature and lack the effective capacity to generate substantial amounts of hydrocarbons. This maturity trend reasonably matches the data of Cruz et al. (1999) in a recent evaluation involving several wells of the Al Sur de la Dorsal petroleum district, eastwards of the study area. Discussion Origin and duration of marl limestone cycles Most of the marl-limestone cycles result from the combination of low siliciclastic supply and abundant carbonate supply derived from high

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Fig. 9. (a) 613C and 618O variation in the limestone and marl sequence of the Los Catutos section (stratigraphic position of the samples at scale). Discontinuous trace indicates that some interstratified limestone-marl couplets were not analysed, (b) 8 3C and 618O as in (a), but samples are represented equidistant to one an other to avoid symbol overlapping. Blank triangles and solid diamonds represent limestone and marl, respectively. Numbers 1, 2 and 3 mark the negative §13C excursions.

TITHONIAN LIMESTONE-MARL RHYTHMITE

219

Fig. 10. 613C v. 8}8O plot showing co-variant distribution of the upper 10 samples in Figure 9 (dashed line) and for the whole set of samples (solid line).

carbonate productivity. Under continuous terrigenous flux, carbonate productivity varies in order to effectively produce the limestone-marl cycles (Einsele & Ricken 1991). However, with a steady carbonate supply, the siliciclastic supply must fluctuate to produce the carbonate dilution in the marl levels. Other processes forming limestone-marl successions, like preferential carbonate dissolution at the seafloor, are not likely for the origin of Los Catutos Member because these rocks were deposited well above the lysocline. Carbonate dissolution due to organic matter degradation and CO2 production (Diester Haass 1991) can also be discarded because dissolution of nannoplankton and body fossils is similar in marls and limestones (Scasso & Concheyro 1999). Marl formation might be favoured by diagenetic carbonate dissolution. Diagenesis and/or weathering often enhance the regular parting parallel to the bedding planes, typical of lithographic limestones (Ricken 1986; Bathurst 1987; Ricken & Eder 1991). Terrigenous content varies between 18 and 30% in marls and limestones of the Los Catutos Member. Single limestone beds are thicker than marl beds. The average thickness

of limestones and marls are, respectively, 0.30 and 0.12 m in the lower part of the section, and 0.5 and 0.05 m in the upper part. Bed thickness and carbonate content point to fluctuations of carbonate productivity as the cause of the limestone-marl alternations (Ricken 199 la, p. 171, fig. 2a; see also Einsele & Ricken 1991, fig. 7). However, the presence of very thin marls with abundant seams and evidence for oxidization in the upper part of Los Catutos Section suggest that diagenetic dissolution could also contribute to the formation of these beds. Following the duration of the Middle Tithonian proposed by Gradstein et al. (1995) an average sedimentation rate of about 5 cm ka~ [ (50m Ma'1: Scasso & Concheyro 1999) was calculated for the limestone-marl rhythmite of Los Catutos. This is relatively high for carbonate shelves, where sedimentation rates usually vary from 3 to 50 m Ma~* (Ricken 199Ib). High sedimentation rates showing periodic variations that resulted in limestone-marl couplets, together with a lack of recognizable unconformities in the section, allowed an estimate of the average duration of each of the marl-limestones cycles of Los Catutos Member as 18.2 ka (Scasso & Concheyro 1999). This value is in the range of

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Fig. 11. (a) Cross-plot of S I + S2 Rock-Eval pyrolysis yields v. total organic carbon (TOC). Fields of 'Good Source Rock' and 'Moderate Source Rock' are given for reference and denote the hydrocarbon generation potential of the samples, (b) Van Krevelen diagram showing changes from kerogen type II to II—III, interpreted either as primary/ preservation variations during organic matter deposition or, alternatively, as an indication of a weathering/telogenetic processes on the kerogen. Both illustrations compare the values of two portions of ZAP A-22, corresponding to the dark brown core of the sample (circle symbol) and to its light brown, weathered and outer part (square symbol). Drops of TOC (8%), SI + S2 (45%) and hydrogen index (40%) values are remarkable, paralleled by an increase in oxygen index (37%) from core to surface of the sample (1 cm sampling distance). Consistently, the kerogen concentrate of the core portion contains approximately 25-30% framboidal pyrite, while this amount is limited to 5-10% in the outer portion.

frequencies for Milankovitch precession cycles (e.g. Fischer 1991; Schwarzacher 1991). Thus, the genesis of the rhythmites of Los Catutos Member is probably related to climatic changes controlled by orbital changes of the Earth. Clay minerals and geochemical results: provenance and palaeogeographic implication Although smectites and illite/smectite minerals dominate throughout the studied section, some levels are characterized by the presence of discrete illite and others by kaolinite. Kaolinite is usually scarce, but may be locally dominant. Variations of clay types in marls and limestones are restricted to some particular couplets, but this is not true for the whole sequence as previously thought (Scasso et al 2002). Expandable minerals can be related to pedogenic alteration of volcanic materials in the source areas, produced by temperate seasonal

regimes. Kaolinite episodes are mainly restricted to thick-bedded marl beds in intermediate intervals of the section. A close relationship between kaolinite presence and thickest marl beds may occur as a product of maximum terrigenous supply in warmer and wetter climates, with higher precipitation rates and consequently good drainage and enhanced runoff, as well as stronger weathering and hydrolysis in the continent (Chamley 1989). Analysis of insoluble residue of marls and limestones shows a negative correlation between A12O3 and SiO2 opposite to the positive correlation observed for most lithographic limestones (Bausch et al. 1999). Excess of silica in relation to other typical detrital oxides, such as alumina, may reflect either aeolian supply as pure quartzose grains or intrabasinal biogenic silica (Arthur & Dean 1991). Microscopic studies show unequivocally that variable intrabasinal biogenic supply is responsible for silica excess in Los Catutos rhythmites. Trends of correlation between alumina and clay mineral contents corroborate this observation (Fig. 8). The

TITHONIAN LIMESTONE-MARL RHYTHMITE

SiO2/Al2O3 v. CaCO3 plot (Fig. 12a) shows maxima for limestone samples (richer in biogenie silica) and above average values (3-5) in limestone-marlstone bedding cycles (Arthur & Dean 1991). However, not all the limestone beds are enriched in silica, as minimum values correspond both to limestones and marls. Finally, the usual negative correlation in the SiO2 v. CaCO3 plot (Fig. 12b), together with the range of silica and carbonate amounts, are typical of limestone-marl couplets originated by varying carbonate input in high carbonate content deposits (Ricken 199la). Plotting of Al2O3/TiO2 v. CaCO3 show variable values (Fig. 12c). Highest values are found in limestones, while very low ones are reported for both marls and limestones and can be related to the widespread presence of smectites, some of which may be of volcanogenic provenance (Arthur & Dean 1991). This is supported by the Na2O/K2O ratio (Fig. 12d), which yields values from 0.5 to 2, indicative of volcaniclastic origin for clay minerals (Arthur & Dean 1991). Sedimentary environment During the Tithonian-Berriasian the Neuquen Basin was a partially isolated sea at high sea

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level (Neuquen Gulf), bounded to the west by a volcanic arc of subdued relief and connected to the Pacific Ocean by sea passages (Uliana et al. 1999; Spalletti et al. 2000). According to the palaeogeographic reconstructions (Leanza & Zeiss 1990; Legarreta & Uliana 1996), Zapala was located 100 km north of the SE coast of the Gulf. The laminated mudstones and marls of the Vaca Muerta Formation are rich in organic matter (2-12% of TOC) and might suggest dysaerobic-anaerobic sea-bottom conditions as a result of a stratified water column and positive hydrological balance in the Neuquen Gulf (Legarreta & Uliana 1996). Anoxic conditions in the Neuquen Gulf could also result from the interchange with anoxic waters from the Pacific Ocean, where upwelling and an oxygen minimum zone were probably developed (Spalletti et al. 2000). Pacific upwelling could have affected the Neuquen Gulf if secondary gyres detached from the main oceanic current (Riggs 1984; Snyder et al. 1990; Scasso et al. 2002). Accumulation of the Los Catutos Member took place in shallow-water and quiet bottom conditions, as suggested by the composition and taphonomy of the fauna (Leanza & Zeiss 1990). Radiolaria, foraminifera and nannoplankton

Fig. 12. Scatter plots for limestones (solid squares) and marls (empty squares), (a) SiO2/Al2O3 ratio v. CaCO3. Maximum values (above 5) correspond to limestones rich in biogenic silica, (b) SiO2 v. CaCO3 showing negative correlation indicative of the transitional variation from limestone to marl and dominance of silica from clastic origin, (c) Al2O3/TiO2 v. CaCO3. Higher values of Al2O3/TiO2 are recorded in limestones, whereas the lowermost values are recorded both in limestones and marls, (d) Na2O/K2O ratio v. CaCO3. Note that Na2O/K2O values are mainly between 0.8 and 2.

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point to open-marine conditions at the sea surface. The high content of kidney-shaped sponge spicules (rhaxes) among siliceous microorganisms is typical of shallow-moderately deep Upper Jurassic limestones in Germany (W. Kiessling pers. comm. 2004). Collapsed ammonite chambers (Fig. 5) indicate that ammonite shells were not filled with sediments when they collapsed. Therefore, shell collapse was the result of burial and septa dissolution (Seilacher et al 1976; Maeda & Seilacher 1996), pointing to rapid burial in the absence of bottom currents. Findings of complete and well-preserved callianasid moults at the top of the Los Catutos Member (Aguirre-Urreta & Scasso 1998) confirm the low-energy environment suggested by the ammonites. Alternating anoxic and disoxic conditions are recorded by isolated levels with benthic fauna and finely laminated layers interbedded with massive beds with Chondrites-like bioturbation. Massive biopelmicrites are rich in sponge spicules, phosphatic particles, glauconite, foraminifera and radiolaria, and poor in terrigenous clasts, suggesting an intrabasinal source for this lithology. The laminated limestones often show thin graded beds and maximum contents of phosphatic particles, terrigenous and pellets suggesting sedimentation from suspended plumes. Some laminae are rich in oriented valves (Fig. 6a) as a result of reworking by weak bottom currents. Most valves and micritized particles in laminated limestones probably originated in shallow-water within the photic zone, by the action of endolithic algae, fungi and cyanobacteria (Kobluk & Risk 1911 a, b; Land & Moore 1980; Reid & Macintyre 2000), and were introduced to the final depositional environment by dilute turbidity currents or suspension plumes, as inferred from their occurrence together with normal gradation and lack of erosive surfaces. The organic assemblage is similar to that of the Vaca Muerta Formation in nearby areas of the basin (Cruz et al. 1999), dominantly derived from marine, probably algal-bacterial, organic matter. High correlation between hydrogen index (HI) and total organic carbon (TOC) in Figure 11 indicates marginal oxic-anoxic conditions for laminated limestones, showing preservation of organic sapropelic matter and high sedimentation rates. In summary, the rhythmites were deposited on poorly oxygenated, low-energy sea bottoms, below superficial open sea waters. The sedimentation rates were high due to the high productivity, together with the contribution of suspension plumes carrying material eroded and transported from proximal areas (Barthel et al.

1994; Pittet et al. 2000; Scasso et al. 2002), which is a typical process for a ramp environment. Wide, gently steeping mounds at distal positions of a ramp would result from rapid accumulation of the rhythmites (Spalletti et al. 2000; Scasso et al. 2002). Such mounds are common in the Late Jurassic, the deep water slope and basin, sponge-algal mounds below normal wave base and near the lower limits of the photic zone (James & Bourque 1992), and can also be compared to the sponge spicule-rich 'distal mud mounds' (Tucker & Wright 1990; James & Bourque 1992) that were also described within the Neuquen Basin in slightly younger sediments (Valanginian) by Palma et al. (2000). Time-equivalent subsurface sequences towards the NW (Mitchum & Uliana 1985) also show lensoid clinoforms, with thicker central sections and wedges thinning both towards the deeper parts and the margins of the basin. Thus, the Los Catutos section represents the deposition in mounds in an intermediate position between a well-oxygenated inner ramp and the deep, euxinic basin. Implications of C isotopes It is well known that micritic limestones can preserve the original environmental 513C signal (e.g. Joachimski 1994). 513C averages 0.95%0 in the Los Catutos section, in good agreement with the average of 1%0 (range between 2.05 and — 3.26%c) calculated for normal marine Tithonian sea water (Ditchfield et al. 1994; Podlaha et al 1998). Values of £13C are, indeed, quite stable about 1.4%o (Fig. 9), close to that of the middle Tithonian ocean (about 1.5%o: Weissert & Mohr 1996), punctuated by three excursions towards negative values (identified as 1, 2 and 3 in Fig. 9). 'Negative excursions' were only considered when they included at least five deviated values. The uppermost excursion shows a covariance between C and O (Fig. 10), which suggests, together with the low 518O, a diagenetic origin for that negative deviation (Jenkyns & Clayton 1986). Presently, 613C decreases in ocean waters as depth increases, as a result of surface photosynthesis and bottom degradation of organic matter. This could have happened during Cretaceous times (Zachos et al. 1989) up to the K/T (Cretaceous-Tertiary) boundary productivity crisis, when productivity decreased due to lower photosynthesis rates, and surface sea water became enriched in 12C, as did the carbonates formed thereafter. Similar productivity crisis, related to massive extinctions in the ocean during periods lacking major polar ice

TITHONIAN LIMESTONE-MARL RHYTHMITE

sheets, were proposed for other boundaries like the Palaeocene-Eocene boundary (Corfield 1994). Lack of great polar icecaps was also proposed for the Late Jurassic (Valdes & Sellwood 1992; Ditchfield et al 1994; Podlaha et al 1998). Weissert & Mohr (1996) explained fluctuations of 513C during the late Jurassic and early Cretaceous as a result of the volume of organic C (drawn out of sea water), buried as organic matter or carbonate. Relevant positive excursions of 813C would be the result of high organic matter burial in the context of monzonic regimes in land and eutrophic conditions in the ocean, whereas returning to negative values would be the result of extensive carbonate sedimentation on the shelves. Alternatively, the existence of a stratified sea with an upper, oxic layer with high productivity and relatively high 613C, and a lower anoxic layer with decaying organic matter, sulphate reduction processes and lower 613C, was proposed by Gruszczyfiski (1998) for part of the Jurassic. Temporary mixing of both layers would lead to decrease of 613C in the ocean surface and the formation of 13 C-depleted carbonates. Because of the relatively short period of time represented in the Los Catutos section, little comparison can be made with other published curves for the Jurassic (e.g. Weissert & Mohr 1996; Gruszczyfiski 1998). In addition, their negative excursions last much longer than those shown in this paper. Carbonate shells from Los Catutos were mainly formed close to the surface of the sea. Minimum values of 613C appear similar in the limestones and marls, and no petrological evidence of decreasing salinity in the sea water was found (cf. Joachimski 1994). Therefore, an estuarine circulation system with a significant freshwater supply seems to be unlikely. Although such a system would have produced anoxic conditions and a decrease in 513C, marine species would have been restricted by the high amounts of freshwater. Moreover, a synchronous decrease in 618O should have been recorded. Alternatively, the negative excursions of 613C can reasonably be related to a productivity crisis in the system or to a temporary mixing of the layers in a stratified sea. The later explanation is more likely, taking into account the relatively short period involved and the lack of biostratigraphical and sedimentological evidence for biological crises. Regarding the elapsed time for the deposition of Los Catutos, about 1 Ma (Scasso & Concheyro 1999), the 8 C excursions can be tentatively related to a E3 periodical variation of the ecosystem (400 ka, Fischer 1991) as the result of

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changes in the eccentricity of the Earth orbit. These considerations are necessarily preliminary, taking into account the limited thickness of the studied unit. Cements, 818O isotopic composition, organic matter and diagenesis of limestones Micritic envelopes around calcareous particles point to diagenetic processes in the marine phreatic zone (Longman 1980). Inclusion-free syntaxial overgrowths are also interpreted as marine phreatic cements (Meyers & Lohmann 1978; Maliva 1989), although they could have originated in other environments. Drusy spar fillings in empty chambers of fossils and undeformed radiolarians tests replaced by calcite suggest carbonate-rich solutions during the early diagenesis. Calcitization of gastropod and some bivalve shells was selective and apparently controlled by the aragonitic or aragonitic-calcitic composition of the shell, which often preserves the original microstructure. Paucity of syntaxial overgrowths on echinoderm spines and plates, together with the scarce lenses of neomorphic blocky spar in the matrix, and the selective calcitization of unstable bioclasts point to local dissolution and allow the influx of freshwater or acid fluids during phreatic diagenesis to be discarded. Blocky spar cements result from marine diagenesis (Freeman-Lynde et al. 1986) and, together with neomorphism and bioclast calcitization, may suggest a later origin under moderate burial and higher temperature (Tucker & Wright 1990), but lack of baroque calcite, sutured stylolites and other features of strong compaction suggest a lack of extreme burial diagenesis. Iron oxides filling the seams could have been derived from the pyrite alteration and, together with some veins filled with blocky sparite, can be assigned to meteoric processes. Dolomite origin is uncertain because its geochemical and isotopical compositions are unknown. In the matrix of the studied sections, the dolomite occurs as very scarce and isolated euhedral rhombs. These features, together with the abundant evidence of marine diagenesis, lack of signs of freshwater phreatic or meteoric diagenesis and the facies context, suggest dolomite formation under marine diagenetic conditions (Friedman 1964, 1989; Zenger & Dunham 1980; Jorgensen 1983; Kastner 1984; Burns & Baker 1987; Lumsden 1988; Graber & Lohmann 1989; All 1995; Swart and Melim 2000- Yoo et al. 2000, among others). 5^O varies from -3.5 to -5.5%o, and S13C from -1 to 1.6%c in the Los Catutos beds.

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These values are well within the 'normal marine limestone field' (Hudson 1977) and the range is similar to that of some Tethyan Upper Jurassic limestones, like Weltenburg, Cerin and Canjuers, but different to some other lithographic limestones, like those of Montardier, Schernfedt, Montsec and Wintershof (Bausch & Joachimski 1994; Bausch et al 1995, 1996). 818O of Tithonian sea water is believed to oscillate about a value of — 1.0%c, ranging from +2.30 to -3.39%, as stated by Ditchfield et al (1994), Podlaha et al (1998) and Veizer et al 1999). Therefore, 518O in Los Catutos limestones is too low to be interpreted as the record of local or global variations in the composition. Hudson (1977) suggested that oxygen isotopic composition of carbonate in epicontinental marine limestones is changed to lighter values through diagenesis; this is not true for carbon, which tends to be more stable (e.g. Veizer et al 1999). Although increasing temperature during burial diagenesis is not the only reason for the negative shifting of 518O (e.g. Mozley & Burns 1993), it is the most likely process in Los Catutos rocks, regarding the distribution of values in the 8ISO-813C plot (Fig. 10). A co-variant distribution is apparent, in spite of the wide dispersion of the points in the graph. Co-variant distribution of C and O (Given & Lohmann 1984; Jenkyns & Clayton 1986; Bausch et al 1996) is normally assigned to diagenetic modification of the original composition of the carbonates. This points to a diagenetic (probably early diagenetic) modification of the original, bulk-rock isotopic composition of the Los Catutos limestones, via organic matter degradation, formation of CO2, carbonate neomorphism and/or precipitation of newly formed calcite cements. This type of covariance is clearly expressed in the upper 10 samples of the section in Figure 9 and is consistent with petrographical evidence of several generations of calcite cements in a closed marine diagenetic system. It is therefore logical to assume that the original, pre-diagenetic, isotopic composition in Los Catutos was close to the maximum (less depleted) values of 618O. An estimation of diagenetic palaeotemperatures assuming an isotopic composition of diagenetic fluids similar to that of the Tithonian sea water (8w = -1.0% SMOW, Ditchfield et al 1994) and an equilibrium fractioning for the oxygen isotopes expressed by T= 16.9 - 4.2l(8c - 8w) + 0.14 (8c - 8w)2

(Irwin et al 1977)

with 8c = <518O (PDB) and 8w = 518O(SMOW), point to palaeotemperatures no higher than 43°C. The palaeotemperature was calculated from a mixture of diagenetic and original components (neomorphized or not), and is therefore a minimum temperature. However, it is in quite good agreement with the thermal history of the unit (Fig. 13a) that suggests a maximum peak temperature of about 65 °C. Temperature estimates from 8ISO for the Los Catutos limestones therefore coincides with data from Rock-Eval pyrolysis, vitrinite reflectivity and fluorescence data, which indicate that limestones and marls are immature and lack the effective capacity to generate substantial amounts of hydrocarbons. Thermal maturity is low (/to = 0.41-0.48) pointing to little burial, in reasonable agreement with the results of Cruz et al (1999) for the area of the Dorsal Neuquina (Huincul High), where the Vaca Muerta Formation is covered by 600-1600m of sediments and shows a vitrinite Ro of about 0.6-0.7%.

Burial history Local burial history for the Zapala-Los Catutos region was modelled on the basis of the reconstructed local and regional stratigraphic column (Table 1; Fig. 13a). This area is located about 60 km to the north of the Dorsal Neuquina, a complex fault structure that underwent several strike-slip movements during the Cretaceous (Ploszkiewicz et al. 1984). Ammonite deformation points to NNW-SSE compression that probably took place at the beginning of the Cenomanian uplift, when the sequence was at its maximum burial, under the influence of the N-S-directed thrusting in the Dorsal Neuquina. Afterwards, the Zapala-Los Catutos area was reached by the Andean east-directed thrusting during the Cenozoic. Therefore, the local geological history was strongly influenced by several events of uplift, first related to the Dorsal Neuquina evolution and then to the Andean orogeny, and the complex structural style in the area results from the interference of these two main structural styles. Los Catutos rocks are believed to undergo initial burial during the Tithonian and up to the Lower Cretaceous. Then a short period of uplift (Intravalaginian tectonic phase) is followed by the deposition of the Agrio Formation. Renewed uplift during the Cenomanian produced

TITHONIAN LIMESTONE-MARL RHYTHMITE

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Fig. 13. (a) One-dimensional basin modelling showing the burial history and temperatures for the Los Catutos Member and the rest of the units in a fictitious well located at Los Catutos. Note that the maximum temperature of burial for Los Catutos Member is about 65 °C. Rocks are thermally immature, (b) Thermal calibration with one control point (square) representing the measured values of vitrinite reflectivity (Ro = 0.41-0.48) in the Los Catutos Member.

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unroofing of the Cretaceous units, like in Neuquen Dorsal to the SE. Accumulation of the Neuquen Group probably provided a thick overburden to the column. Further MaastrichtianTertiary accumulation was almost negligible due to the marginal position of the area with respect to the Maastrichtian-Paleocene sea. Pliocene thrusting from the west resulted in the final uplift of the Jurassic beds. Thermal maturation modelled on the basis of the burial history reconstruction and regional thermal regime is in good agreement with measured vitrinite reflectance (Ro%) data for Los Catutos rocks (Fig. 13b), showing consistency with other palaeotemperature and diagenetic indicators.

Conclusions The lithological and chemical composition of limestones and marls from the Los Catutos Member are similar to those of Tethyan Upper Jurassic sequences in the northern hemisphere. Siliciclastic material made up 18-30% of these rocks. Fluctuations in siliciclastic material probably originated from variations in carbonate productivity. Alumina content proved to be a good indicator of the clastic content of the rocks. However, silica is not reliable as a clastic marker, because maximum amounts of silica in the limestones resulted from high biogenic productivity of siliceous organisms. Smectites and mixed layers dominate the clastic fraction. This, together with the chemical composition of the limestones and marls, point to volcaniclastic rocks exposed to soil processes on land as the source of clastic material. Thick marl beds rich in kaolinite point to a wetter climate, and enhanced hydrolysis and runoff from the continent. Isotopic composition of C in bulk rocks yielded values consistent with those of precipitation from Tithonian sea waters, and show negative excursions of 513C. The negative excursions of 513C may reflect productivity crisis or temporary mixing of water layers in a stratified sea, with a 400 ka periodicity. This fits with the periodicity of the E3 eccentricity change for the Earth's orbit. Rhythmites were deposited under low-energy and poorly oxygenated open-sea conditions, with high sedimentation rates due to high productivity and abundant supply of terrigenous sediments from suspension plumes. The general geometry of the limestone body, together with the palaeoenvironmental conditions, suggest deposition in wide, low-relief mounds with

gentle slopes, at the more distal part of a submarine ramp. The isotopic composition of O in bulk rocks is lighter than that observed in reference curves for Tithonian sea waters. This fact, together with petrographic evidence, point to a modification of the original oxygen isotope signal during limestone diagenesis. Palaeotemperature calculations from 518O are consistent with measured vitrinite Ro and the immature character of the organic matter in the limestones. All data are consistent with the results drawn from ID basin modelling for Los Catutos limestones on the basis of a geological history of burial during the Tithonian and up to the Lower Cretaceous (with a short period of uplift, the Intavalanginian tectonic phase), renewed uplift during the Cenomanian and a significant Late Cretaceous accumulation. Research work was granted by Antorchas Foundation. We thank W. Bausch, M. Joachimski and B. Wenzel for stable isotope analysis in the University of Erlangen, Loma Negra Company for allowing our work in the quarries, H. Lippai, A. Giusiano, L. Castro and P. Bosch for their help in the field work, and S. Quenardelle and P. Leal for their help with the photomicrographs.

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Hemipelagic record of orbitally-induced dilution cycles in Lower Cretaceous sediments of the Neuquen Basin GUILLERMINA SAGASTI Centra de Investigaciones Geologicas, CONICET - UNLP, Calle 1 No. 644, B1900TAC La Plata, Buenos Aires, Argentina Present address: Repsol YPF, Departamento de Geociencias, Paseo de la Castellana 280, 28046 Madrid, Spain (e-mail: gsagastic @ repsolypf. com) Abstract: The lower Valanginian-lower Barremian Agrio Formation in the Mendoza region of the Neuquen Basin (Argentina) accumulated in a hemipelagic outer ramp/basin setting. The formation records the alternation of periods of fine-grained carbonate accumulation and periods of fine-grained clastic deposition, which resulted in a strongly rhythmic succession in which dark clastic hemicycles alternate with light carbonate hemicycles. Clastic and carbonate hemicycles show distinct geochemical signatures that reflect cyclic fluctuations in terrigenous influx to the basin. Clastic hemicycles have relatively high percentages of Fe, Al and K, and less than 60% CaCO3, whereas carbonate hemicycles contain relatively low quantities of detrital elements (Fe, Al and K), and more than 60% CaCO3. Fluctuations in terrigenous supply to the basin occurred in response to the alternation between two contrasting climate regimes and resulted in the formation of dilution cycles. Warm temperate (winter-wet) climate conditions led to the accumulation of clastic hemicycles, whereas arid conditions promoted the deposition of carbonate hemicycles. The ultimate mechanism that controlled the alternation of the dark and light intervals was orbital forcing, as the latitudinal shifting of climate zones was driven by the c. 21 000 years precessional cycle.

Sediments deposited in pelagic and hemipelagic environments commonly show prominent small-scale cyclicity in which carbonate-poor lithologies, such as shales or marls, alternate at regular intervals with carbonate-rich lithologies, such as marly limestones or micritic limestones. The pairing of carbonate-poor and carbonate-rich beds is referred as a 'bedding couplet' (Fischer & Schwarzacher 1984; Einsele & Ricken 1991). The origin of pelagic and hemipelagic rhythmic successions has been a topic of debate for more than 100 years, and is still a subject of discussion. Gilbert (1895) first recognized these rhythms in the Cretaceous Western Interior Basin of the USA, and sought their origin in climate response to forcing by the Earth's precessional cycle of approximately 21 000 years. However, orbital cycles remained unproven in the stratigraphic record until 1976, when Hays et al. (1976) demonstrated regular cyclicity with Milankovitch periodicities in oxygen isotope and other data from deep-sea cores. An alternative explanation for such rhythmically structured successions was proposed by Sujkowski (1958), who suggested a diagenetic

origin based on the redistribution of carbonate from an original uniform marly sequence. Purely or dominant diagenetic origin of limestone-marl alternations has also been postulated by several other authors (e.g. Hallam 1986; Eder 1982; Ricken 1986; Ricken & Eder 1991); however, most workers currently reject diagenesis as the exclusive mechanism for the generation of rhythmic successions. Instead, it is more accepted that carbonate generally moves from clay-rich to carbonate-rich strata, and thus tends to enhance pre-existing bedding. In relation to this, Bohm et al (2003) proposed that rhythmic alternations form through the interaction of an external trigger and diagenetic self-organization. These authors referred to sediment composition, initial porosity or organic content as possible external triggers, yet emphasized that these variables do not necessarily have significance in terms of climate or other rhythmic changes of the environment. Einsele & Ricken (1991) suggested that marllimestone and shale-limestone alternations form mostly in response to variations in carbonate productivity, carbonate dissolution and terrigenous

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 231-250. 0305-8719/05/$15.00 © The Geological Society of London 2005.

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dilution. Many authors concluded that these processes are related to orbitally driven climate changes (Einsele & Ricken 1991; Oloriz et al 1992; Lamy et al. 1998; Sageman et al. 1998, among others). Variations in carbonate productivity are particularly important in the formation of bedding cycles that have a carbonate fraction almost entirely composed of pelagic material and show insignificant signs of dissolution or terrigenous dilution (Einsele & Ricken 1991). Productivity cycles are characterized by variations in the influx/productivity of pelagic carbonate superimposed on a relatively constant influx of terrigenous material, and ideally they produce cyclic sequences in which carbonate hemicycles are thicker than clastic hemicycles (Einsele & Ricken 1991). Dissolution cycles form in response to carbonate dissolution and their development is related to environments affected by fluctuations in the lysocline and carbonate compensation depth (CCD). According to Arthur et al. (1985) the CCD during the Middle Cretaceous was located at water depths of 2-3 km, whereas in the Upper Cretaceous and Tertiary positions ranging from 3.5 to 5 km were reported by Seibold & Berger (1982). Dissolution can also occur in sites above the lysocline particularly if the sediments contain a considerable amount of organic matter (Diester-Haass 1991). Like productivity cycles, these generate rhythmic sequences in which carbonate hemicycles are thicker than clastic hemicycles (Einsele & Ricken 1991). The third mechanism capable of generating clastic-carbonate cycles is clastic dilution. In carbonate environments subject to terrigenous influx, such as epicontinental seas and distal zones of platforms and ramps, periodic fluctuations in terrigenous input become the most important factor that controls the formation of rhythmic alternations of clay-rich and carbonate-rich beds. Terrigenous contribution derives from fluvial, glacial and/or aeolian input, and depends on a complex interplay of geographic and climatic factors. The bedding pattern of dilution cycles is different from that of productivity and dissolution cycles, as clastic hemicycles are frequently thicker than their calcareous counterparts (cf. Einsele & Ricken 1991). Orbitally induced rhythmic sequences have been reported in stratigraphic successions all through the Phanerozoic (Hays et al. 1976; de Boer & Smith 1994; Lamy et al 1998), and numerous examples have been described in Cretaceous sequences, especially from North

America and Europe (Kauffman 1977; Pratt et al. 1985; Fischer et al. 1991; Sageman et al 1997, 1998; Mayer & Appel 1999; Mutterlose & Ruffell 1999; Kobler et al. 2001). However, studies on Cretaceous series from the southern hemisphere, and particularly from South America are rare (cf. Boyd et al 1994). The Lower Cretaceous Agrio Formation in the Mendoza region of the Neuquen Basin constitutes an excellent example of a rhythmic succession. Work conducted by Spalletti et al (1990) documents the first reference to Milankovitch cycles in the sedimentary succession of the Neuquen Basin. More recently, Sagasti (2000, 2002) and Spalletti et al (2001) documented high-frequency cycles in the Lower and Upper Members of the Agrio Formation in the Mendoza region of the Neuquen Basin, and related them to orbitally driven climatic changes. Sagasti (2000, 2002) completed spectral analyses and direct estimations of the frequency of cycles based on their thickness and average sedimentation rate. The author observed periodicities that match the 19-21 ka precession, 41 -53 ka obliquity, 99-127 ka low-range eccentricity and 400 ka high-range eccentricity values (Fischer 1980; de Boer 1982; Arthur et al. 1984). The main objectives of this contribution are: (1) to present the sedimentological and geochemical data of the bedding cycles of the Lower and Upper Members of the Agrio Formation; (2) to discuss the potential factors that controlled their deposition (i.e. carbonate productivity, carbonate dissolution, and clastic dilution); (3) to explore the interaction between sediment supply and climate at the boundary between two climate zones; and (4) to present a comprehensive model for the origin of the cyclicity.

Geological and stratigraphic background The Neuquen Basin is located in central-western Argentina at the eastern foothills of the Andean Range (Fig. 1). Currently, two major structural domains are recognized: the deformed Andean sector to the west; and the submerged Neuquen Embayment to the east (Fig. 1). During the Jurassic and Lower Cretaceous the Neuquen Basin functioned as back-arc basin partially connected to the proto-Pacific Ocean, and was flanked by cratonic areas to the NE (Sierra Pintada System) and SE (North Patagonian Massif), and by the Andean magmatic arc to the west. Sedimentation in the basin was triggered by the inception of a Late Triassic rift system. Post-rift thermal subsidence commenced in

LOWER CRETACEOUS HEMIPELAGIC CYCLES

Fig. 1. Location of the Neuquen Basin indicating present-day structural domains, geographic regions differentiated in this paper (i.e. Mendoza and Neuquen regions) and the position of the studied sections. CIE, Arroyo Cienaguitas; ALO, Arroyo Loncoche; CCH, Cuesta del Chihuido; CL, Canada de Leiva; CR, Rio Seco Cinta Roja; EP, El Porton.

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the Early Jurassic and continued until the Middle Cretaceous (Cobbold & Rossello 2003). The thermal subsidence regime was characterized by a low and constant rate of subsidence (Mitchum & Uliana 1985), punctuated by a number of tectonic-inversion pulses and short intervals of tectonic-induced subsidence (cf. Uliana & Legarreta 1993; Vergani et al 1995). Lower Jurassic-lower Cretaceous post-rift deposits include a thick series of siliciclastic sediments, carbonates and evaporites accumulated in continental and marine settings, which can be subdivided into transgressive-regressive cycles at a variety of scales (Groeber 1946; Mitchum & Uliana 1985; Legarreta & Gulisano 1989; Legarreta & Uliana 1991). The Mendocian cycle (formally Mendoza Group: Stipanicic 1969) spans the Tithonian-Lower Barremian, and comprises a succession of shales, sandstones and carbonates that are distributed over the entire basin. The Agrio Formation is the uppermost unit of the Mendoza Group (Fig. 2). It overlies the Mulichinco and Chachao formations, and is overlain by the Huitrin Formation, both contacts being concordant (Fig. 2). The Agrio Formation (late early Valanginianearly Barremian: Aguirre-Urreta & Rawson 1997) comprises the Lower, Middle (Avile) and Upper members (Fig. 2) (Weaver 1931), and consist of a variety of rock types including siliciclastic and skeletal sandstones, sandy limestones, micritic limestones, marls and shales (Digregorio 1972; Leanza et al 1977; Legarreta et al 1981). Deposition of the unit began with a marine incursion following the major regression of the earliest Valanginian (Legarreta & Gulisano 1989; Legarreta & Uliana 1991). Most of the succession (i.e. the Lower and Upper members) accumulated on a slowly subsiding ramp setting during transgression and highstand of sea level (Legarreta & Gulisano 1989; Legarreta & Uliana 1991). However, these conditions were interrupted by a short episode of sea-level lowstand that caused an abrupt basinward shift of the depositional systems. During this time fluvial, aeolian and marginal marine shales, and sandstones of the Avile Member were deposited in the centre of the basin (Legarreta & Gulisano 1989; Veigaefa/. 2002). The Agrio Formation is widely exposed along the deformed Andean sector of the Neuquen Basin. Overall, the Lower and Upper members include deposits accumulated in two main depositional settings (Leanza et al 1977; Legarreta & Gulisano 1989; Spalletti 1992): (1) a shallow-marine inner-middle ramp setting characterized by siliciclastic-dominated depositional systems in the Neuquen region, and by

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Fig. 2. General stratigraphy and schematic lithology distribution of the Tithonian-early Barremian interval in the Neuquen Embayment, and the Mendoza and Neuquen regions of the Neuquen Basin.

carbonate-dominated depositional systems in the Mendoza region; and (2) a deeper marine outerramp-basin realm that was subjected to predominantly fine-grained siliciclastic (black shale) accumulation in the Neuquen region, and by fine-grained carbonate and clastic sedimentation in the Mendoza region (Fig. 3). Coeval proximal facies representing nearshore and fluvial environments are the Centenario Formation, in the subsurface of the Neuquen Embayment, and the La Amarga Formation, in the southernmost region of the Neuquen Basin (Leanza 1993; Cabaleiro et al 2002). Study area and methodology The area of study is located in the Mendoza region of the Neuquen Basin along an

approximately 200 km-long, N-S-trending corridor, extending from Rio Salado to Rio Colorado (Fig. 1). The outcrops occur in a series of N-S-oriented anticlines that form part of the southern external sector of the Malargiie Fold and Trust Belt (Kozlowski et al 1993), and the Tril Anticline (Kozlowski et al. 1996). In this area the thickness of the Agrio Formation ranges from 200 to 600 m (Fig. 4). The succession is thicker in the El Porton, Rio Seco Cinta Roja, Canada de Leiva and Arroyo Cienaguitas sections, where the Lower, Avile and Upper members are present. In Cuesta del Chihuido and Arroyo Loncoche the Avile Member is absent and the thickness of the unit decreases to about 210 m. The Agrio Formation was logged at detailed scale (1:100), and the Lower and Upper

LOWER CRETACEOUS HEMIPELAGIC CYCLES

235

members (which are the focus of this study) were sampled for petrographic and inorganic geochemical analyses. The facies were initially defined in outcrops based on their texture, primary sedimentary structures and colour. Subsequent petrographic characterization allowed a more precise definition of the microfacies (Sagasti 2002). Samples from four representative localities were analysed for their content of major and trace elements, CaCO3 and insoluble residue (Table 1). One gram of sample was ground and dissolved with HC1 (0.5 N). In the soluble portion, the contents of Ca2+, Fe (Fe2+ and Fe3+), A13+, Mg2+, Mn2+, K+ and Na+ were determined by atomic absorption spectrometry. Measurements were made on a Metrolab spectrophotometer, Model 250 AA using monocathodic lamps. The flame was air-H2C2 (acetylene), except for the determination of A13+, which was N2O-H2C2. Carbon dioxide was determined by volumetry; insoluble residue was determined by gravimetry. All geochemical analyses were conducted in the geochemistry laboratory of the Centro de Investigaciones Geologicas (CONICET-UNLP).

The Lower and Upper members of the Agrio Formation: facies and geochemistry of the bedding cycles

Fig. 3. Schematic distribution of facies and depositional environments during accumulation of the Lower and Upper members of the Agrio Formation, and coeval deposits of Centenario and La Amarga formations.

The Lower and Upper members of the Agrio Formation (referred to here as Lower Agrio and Upper Agrio) comprise intercalations of shales, marls and micritic limestones. The most striking attribute of the succession is its rhythmic bedding, defined on the lithological alternation of a dark-coloured carbonate-poor end member or clastic hemicycle, and a light-coloured carbonate-rich end member or carbonate hemicycle (Fig. 5). A cross-section through the study area reveals a gradual decrease in the proportion of fine-grained siliciclastic v. fine-grained carbonate intervals from south to north (Fig. 4). The rhythmic character of the unit also improves towards the north. Visual identification of the cycles in outcrops is facilitated by variations in bed colour, bed thickness and induration/weathering profile of the bedding cycles (the former attribute is mainly a function of the CaCO3 content). Diagenesis has caused only a minor post-depositional overprint (Sagasti 2002); as a result, the primary depositional character of this rhythmic succession has been exceptionally preserved. Clastic hemicycles (CaCO3 <60%) include marls and dark calcareous shales (Fig. 5). Marls

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Fig. 4. Stratigraphic correlation of the Agrio Formation in the Mendoza region. Note the decrease in fine-grained clastic intervals from south to north.

are medium-dark grey and commonly show massive, thickly-laminated and slightly bioturbated fabrics. Faunal content is scarce, and comprises bivalves, echinoderms, gastropods, benthic foraminifers and ammonites. Dark shales display mm-scale lamination and very little to no bioturbation. Faunal content is very scarce and includes thin-shelled pelecypods, benthic foraminifers and ammonites. Clastic hemicycles are 0.05-5.65 m thick; their average thickness is 0.87 m. Carbonate hemicycles (CaCO3> 60%) comprise grey-light-blue, tabular-bedded, massive bioturbated micritic and marly limestones (Fig. 5). Trace fossils of Thalassinoides are common (Fig. 6a, b), and a few beds also display Planolites, Paleophycus and/or Teichichnus (Sagasti & Poire 1998). Petrographic study reveals the supremacy of mudstone/wackestone textures, with a minimal contribution of packstone/wackestone facies (Fig. 6c-g). Faunal content is dominated by microfossils, with radiolarians, sponge spicules, and benthic

foraminifers being common. Bivalves (Lucinidae and Inoceramus) and ammonites are frequent, whereas echinoderms, gastropods and serpulids are sparse. Carbonate-rich units contain a variable proportion of framboidal pyrite, finely disseminated through the matrix. In addition, subhedral-euhedral pyrite occurs as a fossil replacement. Carbonate hemicycles tend to be thinner than the corresponding clastic hemicycles; their thickness varies between 0.03 and 1.20 m, and the average value is 0.24 m. Bedding couplets or cycles were defined from the base of the clastic hemicycle to the top of the overlying carbonate hemicycle (Fig. 5). The average thickness of the bedding couplets is 1.40 m (n = 103). Variations in the thickness of the bedding cycles are especially controlled by fluctuations in the thickness of the clastic hemicycles, as revealed by the excellent correlation between the thickness of the clastic hemicycles and that of the bedding couplets (Fig. 7). Based on sedimentological and fossil attributes, the cycles are interpreted to represent

237

LOWER CRETACEOUS HEMIPELAGIC CYCLES Table 1. Chemical composition of representative samples from the Lower and Upper Agrio at Arroyo Loncoche, Arroyo Cienaguitas, Canada de Leiva and Rio Seco Cinta Roja IR(%)

Fe203(%)

A12O3(%)

K2O(%)

Na2O(%)

MgO(%)

Mn (ppm)

Arroyo Loncoche alo 1 72.34 alo 2 76.45 alo 3 76.51 alo 4 92.37 alo 5 76.46 alo 6 90.96 alo 7 92.40 alo 8 93.92

22.30 14.80 17.40 4.35 18.20 6.35 4.60 2.75

0.59 0.21 0.26 0.18 0.18 0.43 0.15 0.36

0.46 0.39 0.31 0.04 0.29 0.13 0.15 0.12

0.038 0.115 0.134 0.030 0.113 0.038 0.057 0.048

0.084 0.054 0.057 0.028 0.055 0.035 0.032 0.035

0.80 0.90 0.80 0.60 0.83 0.97 1.00 0.97

811 632 653 291 211 242 252 491

Arroyo Cienaguitas cie 1 68.28 cie 2 65.02 cie 3 47.15 cie 4 79.31 cie 5 90.52 cie 6 41.14 cie 7 78.57 cie 8 61.60 cie 9 90.10 cie 10 39.29 cie 11 53.05 cie 12 39.73 cie 13 59.46

28.15 26.35 44.15 16.90 7.20 49.05 10.85 30.75 6.00 46.10 30.00 47.55 17.15

0.51 0.76 1.29 0.58 0.25 1.33 0.94 1.57 0.83 2.20 1.92 2.60 4.16

0.20 0.31 0.52 0.36 0.06 1.13 0.47 0.78 0.06 1.13 0.85 1.47 0.94

0.048 0.108 0.168 0.108 0.035 0.306 0.100 0.302 0.038 0.250 0.220 0.260 0.180

0.063 0.042 0.042 0.039 0.043 0.051 0.049 0.050 0.049 0.050 0.052 0.044 0.042

0.72 1.16 0.57 0.97 1.16 1.40 0.55 1.03 0.90 0.43 0.43 0.47 6.83

341 442 173 332

1162

Canada de Leiva cl 1 45.99 cl 2 36.97 cl 3 58.03 c!4 91.91 cl 5 89.33 cl 6 89.29 cl 7 74.50 cl 8 46.64 cl 9 87.50 cl 10 42.14 cl 1 1 69.64 cl 12 35.12 cl 13 86.85

44.50 52.45 35.00 4.00 6.90 4.10 16.75 44.10 6.65 44.40 18.95 49.15 7.60

1.78 1.49 0.76 0.58 0.42 0.31 1.33 1.62 0.93 2.60 1.46 1.97 0.79

0.64 0.80 0.70 0.11 0.29 0.23 0.56 2.46 0.38 1.79 1.19 1.04 0.28

0.170 0.230 0.185 0.050 0.091 0.040 0.160 0.220 0.050 0.190 0.130 0.218 0.085

0.055 0.077 0.054 0.030 0.051 0.023 0.039 0.056 0.028 0.038 0.042 0.049 0.035

0.60 0.80 0.60 0.68 1.33 0.43 1.03 0.38 1.67 0.67 0.47 0.85 1.23

381 305 721 291 542 232 721 620 232 387 310 321 492

Rio Seco Cinta Roja cr 1 38.39 cr 2 74.52 cr3 61.86 cr4 85.71 cr5 82.14 cr6 61.43 cr 7 63.57 cr 8 49.07 cr9 82.14 cr 10 35.73 crll 82.14 cr!2 71.43 cr 13 76.79 cr 14 82.14

34.80 21.00 10.60 6.75 9.85 4.85 26.90 30.80 9.10 43.80 9.25 18.45 10.40 11.50

0.49 0.39 0.30 0.11 0.31 0.19 0.77 2.00 0.47 0.95 0.32 0.81 0.43 0.56

0.79 0.30 0.49 0.17 0.57 0.11 0.75 0.72 0.42 1.25 0.49 0.68 0.51 0.42

0.110 0.020 0.050 0.020 0.070 0.020 0.160 0.160 0.060 0.160 0.050 0.100 0.070 0.090

0.028 0.030 0.026 0.024 0.046 0.028 0.036 0.054 0.033 0.032 0.033 0.036 0.032 0.033

9.33 1.55 0.43 0.48 0.67 0.52 0.72 1.17 0.68 1.75 0.82 0.73 0.73 0.60

232 155 78 232 310 310 232 232 232 232 155 232 232 232

Sample

CaCO3(%)

1641

382 542 361 391 232 232 310

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Fig. 5. Main sedimentary and geochemical features of the bedding cycles of the Lower and Upper Agrio. The alternation of dark, more easily erodable clastic hemicycles and light, resistant carbonate hemicycles facilitates their identification in outcrop.

hemipelagic deposition in a low-energy outer ramp/basin environment. Variations in the degree of oxygenation of the bottom waters are inferred from the colour, faunal diversity and degree of bioturbation (cf. Savrda et al 1991, and references therein). Dark shales and marls reflect suboxic conditions, whereas limestones are interpreted to indicate more oxygenated conditions. Results of organic geochemical analyses from the lower intervals of Lower and Upper Agrio in the centre of the Neuquen Basin (south of the area studied in this paper) from Tyson et al. (2005) support this interpretation. Total organic carbon (TOC) mean values of 2-3% (Lower Agrio) and 0.43% (Upper Agrio) characterize shales and marls with usually less than 60% of CaCO3 (and thus comparable to the clastic hemicycles defined in this paper). TOC values together with the presence of benthic fauna were interpreted by Tyson et al. (2005) as indicative of suboxic conditions.

Analyses of major and trace elements, CaCO3 and insoluble residue (Table 1) show that clastic and carbonate hemicycles are geochemically different. A comparison between the proportion of A12O3, Fe2O3, K2O, Na2O, MgO and Mn, and the percentage of CaCO3 and insoluble residue showed that the percentage of detrital elements (A12O3, Fe2O3 and K2O) is related to the lithology (clastic v. carbonate lithologies), whereas the amount of Na2O, MgO and Mn reveal no clear relationship with either carbonate or clastic hemicycles (Fig. 8). The percentage of detrital elements is higher in clastic hemicycles than in carbonate ones (Fig. 8). The content of A12O3 varies between 0.52 and 2.46% in clastic hemicycles, with a mean of 1.08%; and between 0.04 and 1.19% in carbonate hemicycles, with a mean of 0.37%. The percentage of Fe2O3 ranges from 0.49 to 4.16% in clastic hemicycles, and from 0.11 to 1.57% in their carbonate counterparts. Fe2O3 mean values for clastic and carbonate

LOWER CRETACEOUS HEMIPELAGIC CYCLES

239

Fig. 6. (a) & (b) Aspect of Thalassinoides showing their typical T- and Y-shaped branches (arrows in a and b respectively), (c)-(f) Photomicrographs of micritic limestone microfacies showing mudstone and wackestone textures. Note radiolarians in (c), benthic foraminifers in (d), Inoceramus in (e), and gastropods in (f). (g) Photomicrographs of packstone microfacies with abundant benthic foraminifers. Plane-polarized light: (c), (e) & (f). Cross-polarized light: (d) and (g).

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hemicycles are 1.81 and 0.55%, respectively. Finally, the content of K2O varies between 0.11 and 0.30% in clastic hemicycles and between 0.02 and 0.30% in carbonate hemicycles, while the mean content is 0.20% in clastic hemicycles and 0.08% in carbonate hemicycles. Percentage variations of these components are essentially in phase with fluctuations in the amount of insoluble residue (Fig. 9).

Discussion

Fig. 7. (a) Outcrop photograph showing thickness variations in bedding cycles, (b) Thickness distribution of carbonate and clastic hemicycles compared with bedding couplet thickness in the Arroyo Cienaguitas section. Note the correspondence between the thickness curves of the clastic hemicycles and the bedding couplets.

Orbital forcing and the genesis of the cycles in Lower and Upper Agrio Sedimentological and geochemical features of the Lower and Upper Agrio point to a depositional origin of the cyclicity, and results from periodicity analyses of the cycles yield values that closely match orbital cycles (Sagasti 2000, 2002). In particular, the majority of the bedding couplets represent time intervals of about 20 ka (Fig. 10), and thus can be related to the 19-21 ka precessional signal (Fischer 1980; de Boer 1982; Arthur et al 1984). Milankovitch cycles affect the climate through variations in the degree of insolation. These orbital-scale climate changes induce environmental changes that have an effect on the composition or other properties of the sediments produced in source areas and deposited elsewhere. Climate-driven environmental changes such as growth and decay of ice sheets, latitudinal migration of climatic zones, varying precipitation due to a changing land-sea temperature contrast, switches between monsoonal and zonal wind systems, and changes between vigorous and sluggish oceanic-circulation regimes are potentially recognizable in climatically sensitive depositional settings. Among the susceptible depositional environments are lakes, outer shelf/ramp settings subject to hemipelagic sedimentation, pelagic environments around isolated carbonate platforms, and deep-sea environments between the lysocline and the CCD (cf. Barron et al. 1985; Prell & Kutzbach 1987; Einsele & Ricken 1991). In pelagic and hemipelagic settings orbital forcing is believed to affect carbonate productivity, carbonate dissolution and carbonate dilution by terrigenous sediment (Arthur et al. 1984; R.O.C.C. Group 1986; Arthur & Dean 1991; Einsele & Ricken 1991). As all these variables have the potential to react to orbital cycles, the genuine cause of the cyclicity may be the product of their combined effect. However, under particular conditions, the influence of one

LOWER CRETACEOUS HEMIPELAGIC CYCLES

241

Fig. 8. Relationship between the proportion of detrital and non-detrital component and calcium carbonate in clastic and carbonate hemicycles. Fe2O3, A12O3 and K2O show a negative correlation with calcium carbonate, whereas Na2O, MgO and Mn show no correlation.

specific variable may become strong enough to overcome the effect of the others and, thus, provide an unequivocal signal. In the Lower and Upper Agrio the bulk of the fine-grained carbonate fraction is composed of micrite. Although the precise origin of micrite is often difficult to assess, it is proposed here

that carbonate productivity in the shallower areas, bioerosion and mechanical erosion of shallower ramp carbonates have largely contributed to the amount of micrite deposited in the succession. In contrast, pelagic calcareous material, represented by a poorly diversified association of calcareous nannofossils (Concheyro &

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Fig. 9. In-phase fluctuations in the percentage of detrital components (Fe2O3, A12O3 and K2O) and insoluble residue (IR) in Arroyo Cienaguitas, Canada de Leiva and Rio Seco Cinta Roja sections.

LOWER CRETACEOUS HEMIPELAGIC CYCLES

Fig. 10. Frequency histogram of the bedding couplet duration. The example is from the Upper Agrio in the Arroyo Cienaguitas section. Bedding couplet duration was calculated using the mean sedimentation rate and the thickness of the cycles (cf. Gilbert 1895).

Sagasti 1999), is considered a subsidiary component. Prevalence of micrite over biogenic pelagic carbonate argues against pelagic productivity as the major cause for the formation of these carbonate cycles. However, it does not rule out the potential influence of shallow-water carbonate productivity (e.g. Eberli et al 1997). Estimating the water depth during accumulation of the Lower and Upper Agrio, based on the depositional profile of the basin, helps to evaluate the potential influence of dissolution cycles. Following Hampson (2000), the inclination of the slope in the shoreface zone ranges from 0.1° to 0.3°, the gradient along the inner ramp zone varies from 0.05° to 0.1°, and along the middle ramp from 0.01° to 0.03°. Considering these values and the distribution of depositional environments, as shown in Figure 3, the estimated water depth at the boundary between the middle and outer ramp settings would range between 80 and 200 m. Even though these values are at best a rough calculation, considering the physiographical attributes of the Neuquen Basin they seem reasonable. In this respect, the ramp-type profile of the basin and the development of lowstand-wedge deposits, such as the Avile Member, indicate a shallowbasin setting (Legarreta & Gulisano 1989; Veiga & Vergani 1993). Assuming that the CCD during the Lower Cretaceous was located at 2-3 km (Arthur et al. 1985), and that the water depth in the

243

outer ramp/basin environment during accumulation of the bedding cycles was of the order of 100-200 m, it is concluded here that carbonate dissolution did not contribute to the origin of the cycles. This conclusion is further supported by the relative low organic matter content of the succession. In the Lower and Upper Agrio the sedimentology and bedding pattern of the succession, characterized by bedding couplets with clastic hemicycles commonly thicker than their calcareous counterparts, point to terrigenous dilution as the dominant mechanism for the origin of the cycles. In addition, the geochemical attributes of the bedding couplets strongly support regular changes in terrigenous input between the carbonate and clastic part of the cycles. For example, the percentage of detrital elements (Fe, Al and K) shows a negative correlation with the amount of calcium carbonate, and the quantity of these elements varies in phase with the amount of insoluble residue. In contrast, the proportion of non-detrital elements (Mg, Mn, Na) appears to be independent of the calcium carbonate and insoluble residue content. A similar behaviour was reported by Dean & Arthur (1998) in Early Cretaceous cyclic sequences of the North Atlantic Ocean, and by Boyd et al. (1994) in Late Cretaceous cyclic successions from the Exmouth Plateaux (off NW Australia). In both studies this geochemical behaviour was linked to cyclic fluctuations in terrigenous influx. The geological setting and palaeoclimatic context of the Neuquen Basin (Fig. 11) also favour terrigenous dilution as the key driving mechanism for the origin of cyclicity. First, the basin developed as an interior sea with a ramptype profile partially connected to the protoPacific Ocean (Fig. 1) (Spalletti et al. 1999), in close proximity to three potential source areas: the Andean magmatic arc; the North Patagonian Massif; and the Sierra Pintada System (Fig. 3). Thus, it was probably subject to permanent clastic supply. Secondly, according to existing palaeoclimatic and palaeogeographic reconstructions (cf. Scotese 2000), the Neuquen Basin was situated at the boundary between two climatic zones (Fig. 11), and therefore was prone to experience climate variations, which would ultimately have regulated terrigenous influx to the marine basin. In summary, the sedimentology and bedding pattern of the rhythmic succession, the geochemical character of the cycles, the geological setting, and the palaeoclimate and palaeogeographical context of the Neuquen Basin point to cyclic variations in clastic supply as the prime

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Fig. 11. (a) Palaeogeographic setting of the Neuquen Basin during the Early Cretaceous indicating main source areas for siliciclastic sediments (adapted from Spalletti et al 1999). (b) Location of palaeoclimatic boundaries along Gondwanaland during the Early Cretaceous (from Scotese 2000). Note the position of the North Patagonian Massif and southern area of Neuquen Basin at the border between the arid and warm temperate climatic zones.

controlling mechanism for the rhythmicity observed in the Lower and Upper Agrio. Therefore, the bedding couplets are interpreted as dilution cycles.

Model for Agrio dilution cycles: sedimentsupply fluctuations in response to oscillating climate regimes Sediment supply is controlled by a complex interplay of geographic and climatic factors such as relief, the size and temporary storage capacity of the drainage area, climate, amount and seasonality of precipitation, and vegetation cover (Perlmutler & Matthews 1989; Perlmutler et al. 1998). Fluctuations in sediment supply occur at different orders of hierarchy, in relation to long- to short-term variations in palaeogeography and palaeoclimatic conditions. Milankovitch-scale variations in sediment supply are particularly important in regions located at the boundary between climate regimes, because

these areas can be exposed to regular climate changes (Van der Zwan 2002). Three potential source areas flanked the Neuquen Basin during the Upper Jurassic Lower Cretaceous time: Sierra Pintada System; North Patagonian Massif; and Andean magmatic arc (Fig. 3). The relative influence of these source areas has been a subject of discussion in previous papers, and up to now remains controversial. For example, Uliana et al. (1977) suggested a western source for the Neocomian interval; whereas Mitchum & Uliana (1985) postulated that during that period most of the sediments entered the basin across its SE flank. Alternatively, Legarreta & Uliana (1991) indicated that the arc region to the west played a significant role as sediment source, particularly during sea-level lowstands when basinward shifting of the depositional systems allowed transport of material from the arc to the central parts of the basin. More recently, Eppinger & Rosenfeld (1996) conducted the first study of provenance of the siliciclastic sediments accumulated in the Neuquen Basin during the Upper Jurassic and

LOWER CRETACEOUS HEMIPELAGIC CYCLES

Lower Cretaceous. They agreed with Legarreta & Uliana (1991) on the arc provenance during sea-level lowstands, but also indicated that during highstand of sea level most of the clastic sediment was sourced by the North Patagonian Massif, and only a negligible proportion came from the Sierra Pintada System. Deposition of the Lower and Upper Agrio took place during transgression and highstand conditions of sea level (Legarreta & Gulisano 1989; Legarreta & Uliana 1991). Therefore, under such circumstances the North Patagonian Massif is considered the main active source area (cf. Eppinger & Rosenfeld 1996). SE-NW progradation of clinoforms in the Upper Member of the Agrio Formation (Spalletti et al. 2001), as well as the regional trend towards more carbonate-dominated successions from south to north (Fig. 3), support this interpretation. Moreover, in the Mendoza region a detailed correlation of the distal ramp/basin facies of the Lower and Upper Agrio also reveals a gradual decrease in the proportion of fine-grained siliciclastic v. fine-grained carbonate intervals from south to north, away from the terrigenous sediment source area (Fig. 4). Siliciclastic sediments were supplied to the marine depocentre through fluvial systems developed towards the south and SE, as indicated by the distribution of proximal deposits of the La Amarga and Centenario formations (Fig. 3). Once the sediments reached the marine environment they were reworked and redistributed by waves and currents throughout the basin, and only the finegrained fraction was able to reach the more distal areas of the basin (i.e. the Mendoza region). Based on palaeogeographical and palaeoclimate reconstructions for the Lower Cretaceous, the Neuquen Basin was located between 42° and 50°S latitude (Spalletti et al 1999), at the boundary between two climate regimes (Scotese 2000): (1) arid to the north; and (2) warm temperate with seasonal rainfall (cf. Cuneo 2003) to the south (Fig. 11). This palaeoclimatic context has significant implications for the amount and mode of sediment supply, as the basin was almost certainly exposed to regular climate changes in response to orbital-driven latitudinal shifting of climatic zones (cf. Oglesby & Park 1989). Shifts in the atmospheric circulation systems capable of inducing shifts in the climatic belts have been proved to occur at precessional frequencies during the Late Quaternary of Chile (Lamy et al. 1998). According to these authors, the latitudinal migration of the climatic belts promote significant climate changes and modifications of the terrestrial sedimentary

245

environments, which in turn alter the terrigenous sediment input to the marine basin (Lamy et al. 1998). Furthermore, results from forwardmodelling analyses of sediment supply conducted by Van der Zwan (2002) not only reveal the influence of orbitally induced climate changes on sediment supply, but also demonstrate that the influence of Milankovitch-scale climate variations on terrigenous input is particularly strong during greenhouse periods (such as the Cretaceous), as low-amplitude highfrequency sea-level fluctuations do not obscure the sediment-supply signature (Van der Zwan 2002). Taking into account the above, it is concluded that fluctuations in clastic dilution during accumulation of Lower and Upper Agrio occurred as a consequence of recurring latitudinal migration of climate zones forced by the precessional cycle. The alternation between warm temperate and arid climatic conditions in the North Patagonian Massif and southern part of the Neuquen Basin caused changes in the amount and mode of terrigenous sediment supply, and thus contributed to the formation of clastic and carbonate hemicycles in the Mendoza region. During one phase of the precession cycle the North Patagonian Massif and southern part of the Neuquen Basin experienced warm temperate climate conditions with seasonal rainfall. If the atmospheric circulation during the Lower Cretaceous were comparable to the atmospheric circulation of the Upper Jurassic (cf. Parrish et al. 1982; Barron & Moore 1994), then a plausible explanation for this seasonal rainfall is the change in the direction of the winds during summer and winter. If true, the arrival of cool and humid air masses from the western ocean during winter promoted important pluvial precipitation on the continent, which resulted very effectively in releasing and transporting the erosion products to the basin. High runoff and supply of terrigenous sediments to the marine basin promoted the accumulation of clastic hemicycles (Fig. 12). Conversely, during the opposite phase of the precession cycle, the North Patagonian Massif and southern part of the Neuquen Basin were subject to more arid conditions. The arid climatic regime inhibited chemical weathering and runoff. Under these circumstances carbonate material imported mainly from the carbonate-dominated shallow ramp located to the east (and to a less extent from the south) and generated in situ (i.e. organic pelagic carbonate) became the major source of sediments, favouring the accumulation of carbonate hemicycles (Fig. 12).

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Fig. 12. Model for the origin of clastic-carbonate bedding cycles in the Lower and Upper Agrio. According to this model warm temperate (winter-wet) climate conditions in the North Patagonian Massif and southern area of the Neuquen Basin lead to accumulation of clastic hemicycles in the Mendoza region, whereas arid conditions promoted the formation of carbonate hemicycles. Palaeolatitudes and climate boundary between arid and warm temperate zones are based on Spalletti et al. (1999) and Scotese (2000).

LOWER CRETACEOUS HEMIPELAGIC CYCLES

Conclusions •

•

•

•

•

•

In the Mendoza region of the Neuquen B asin, the Lower and Upper members of the Agrio Formation are characterized by a rhythmic succession defined by the alternation of clastic and carbonate hemicycles. Diagenesis has caused only minor post-depositional overprint, and thus the rhythmic character of the succession is interpreted to be a primary depositional feature. Sedimentological attributes of the cycles indicate hemipelagic deposition in a lowenergy outer ramp/basin environment. Suboxic-oxic conditions were inferred from the colour, faunal diversity and degree of bioturbation of the cycles. Dark shales and marls reflect suboxic conditions, whereas limestones are interpreted to indicate more oxygenated conditions. Clastic and carbonate hemicycles have different geochemical signatures: clastic hemicycles show relatively high percentages of Fe, Al and K, and less than 60% calcium carbonate content; whereas carbonate hemicycles contain relatively low quantities of detrital elements (Fe, Al, and K) and more than 60% calcium carbonate content. These geochemical differences are interpreted to reflect cyclic fluctuations in terrigenous influx. Considering the Sedimentological and geochemical attributes of the cycles in conjunction with the geological and palaeoclimatic setting of the Neuquen Basin, the bedding couplets of the Lower and Upper Agrio are interpreted to have formed in response to recurring variations in clastic supply and therefore they represent dilution cycles. Fluctuations in clastic dilution occurred in response to recurring latitudinal migration of climatic zones forced by the precessional cycle. The alternation between warm temperate (winter-wet) and arid climate conditions in the North Patagonian Massif and southern area of the Neuquen Basin during the Early Cretaceous incited changes in the amount and mode of terrigenous sediment supply, and thus triggered the formation of clastic and carbonate hemicycles. The model for clastic dilution proposed here highlights the importance of palaeogeography and latitudinal shifting of climatic belts in the development of hemipelagic orbital cycles, and may serve as an analogue in the study of hemipelagic rhythmic successions deposited in similar settings worldwide.

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This study is part of the work for the author's PhD, project. Funding for research was provided by the CONICET and by the University of La Plata. The author thanks N. Canessa, R. Lopez and E. Schwarz for their support in the field. She also thanks S. Ballent and A. Concheyro for their valuable help with the micro- and nannofossil studies. The author is grateful to reviewers P. de Boer and A. Ruffell for their essential and constructive comments.

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Organic fades variations in the Valanginian-mid-Hauterivian interval of the Agrio Formation (Chos Malal area, Neuquen, Argentina): local significance and global context RICHARD V. TYSON, PHILIP ESHERWOOD & {CATHERINE A. PATTISON School of Civil Engineering and Geosciences, Drummond Building, University of Newcastle, Newcastle upon Tyne NE1 7RU, UK (e-mail: [email protected]) Abstract: Marine shales and marls of the Valanginian-Hauterivian Agrio Formation have been studied at five localities in order to assess lateral variations over a 100 km S-N, shelf to basin transition. The two main organic-rich intervals at the base of the Pilmatue Member (Valanginian) and the base of Agua de la Mula Member (late early Hauterivian) have been characterized using a combination of bulk organic chemistry and palynofacies. Except for the former at the southern end of the transect, both intervals have mean total organic carbon (TOC) contents of 2-3% and are dominated by marine amorphous organic matter, suggesting a similar dysoxic genetic organic facies. The mean hydrogen indices determined from the slope of S2 v. TOC are 174 in the Pilmatue Member, but 387 in the basal part of the Agua de la Mula Member, a difference that mainly reflects the range in thermal maturity (late v. early oil window, respectively). Significant lateral variation occurs in the Pilmatue Member, with dark organic-rich intervals being rare in the south but dominant at the northern (distal) end of the transect; this trend is matched by a progressive increase in the peak or mean carbonate-free TOC and hydrogen indices, the latter reaching 6% and 297, respectively, near Estancia Pampa Tril. The bulk of the Agua de la Mula Member in the south is developed in organic-poor oxic facies, with a predominance of terrestrial phytoclasts and type IV kerogen, but dysoxic-anoxic conditions apparently predominate in the northern area. Valanginian-Hauterivian black shale facies appear generally rare on a global basis, but their occurrence can be related to the combination of the progressive rise in sea level during the Early Cretaceous and locally more restricted conditions.

The Neocomian Neuquen Basin is a back-arc basin with a NW-SE-trending axis and a predominantly SE source of siliciclastic sediment, a combination producing a transition from fluvial and inner-shelf elastics in the SE to shallow carbonate platform and then low-energy basinal facies in the NW (Leanza 1981; Legaretta & Uliana 1991, pp. 431-432). The basinal facies is best developed in the provinces of central and northern Neuquen and Mendoza (Fig. 1), with the elongate basin depocentre approximately along the line of the Rio Neuquen (Urien & Zambrano 1994; Kozlowski et al 1998). As the Neuquen Basin experienced a relatively simple 'thermal sag' mode of subsidence (Legaretta & Uliana 1991, p. 432) a low-gradient depositional ramp developed in which the nature and extent of the basinal facies were strongly influenced by variations in global sea level (Uliana et al. 1999). The maximum Mesozoic marine flooding of the basin during the late Valanginian and late Hauterivian transgressions resulted in the deposition of the dark-coloured

basinal shales, marls and micritic limestones of the Agrio Formation (Riccardi 1988), which extend over nearly half of the marine part of the basin (Legaretta & Uliana 1991, p. 442). The Agrio Formation includes organic-rich and bituminous 'black shale' facies (Leanza 1981, p. 85; Uliana & Legaretta 1993, p. 397) which, although less extensive than the main Tithonian-Berriasian Vaca Muerta Formation source rock, contributed to the 'Agrio?' petroleum system that has a cumulative production of over 164 million barrels of oil (Urien & Zambrano 1994, p. 530). Because of subsequent shallowing and emergence, the Agrio Formation represents the final development of marine source-rock facies in the Neuquen Basin (Uliana et al. 1999, p. 26) - a distinct contrast to most areas of the world where such facies do not peak until the Aptian-Turonian. This contribution presents new data on the organic facies of the Agrio Formation from five outcrop sections in the Chos Malal region of Neuquen Province (Fig. 1), which together

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 251-266. 0305-8719/05/$15.00 © The Geological Society of London 2005.. 2005.

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Fig. 1. Location map for the Neuquen Province and the sections studied or mentioned in the vicinity of Chos Malal. Pichi Mula and Agua de la Mula are contiguous sections for the upper and lower Agrio Formation, respectively, and thus only the latter is named on the map. Maximum southern and eastern extent of Agua de la Mula Member organicrich basinal facies from Uliana et al. (1999); approximate southern edge of the oil window (0.6% vitrinite reflectance) from Urien & Zambrano (1994).

form an approximately 90-100 km N-S transect. These preliminary results are compared with previous data and observations from outcrop and subsurface studies, with particular emphasis on large-scale stratigraphic and lateral variability. The general global rarity of Valanginian and Hauterivian shelf source-rock facies also gives the Agrio Formation a wider significance, and thus we also briefly assess its relationship to coeval records of marine dysoxia-anoxia in other parts of the world.

Stratigraphy and facies In the basin depocentre the Agrio Formation has a maximum thickness of around 1500-1600m (Riccardi 1988; Urien & Zambrano 1994), but

this is reduced to around 200-300 m in the north (e.g. in the Malargiie area of Mendoza) where there was lower subsidence and less siliciclastic input (Riccardi 1988; Sagasti & Poire 1998). In the Chos Malal study area the thickness is around 1175-1220m (Cruz et al. 1996; Aguirre-Urreta et al 1999). As well as the thickness varying laterally, the base of the formation is also diachronous, varying from latest early Valanginian (curacoensis Zone) in the depocentre near Chos Malal to latest Valanginian (angulatiformis Zone) further to the south at Agua de la Mula and El Salado (Aguirre-Urreta et al. 1999, p. 35). This diachroneity (younging in a proxima direction) reflects the earlier development of the transgressive basinal facies in the deeper depocentre, and its progressive onlap onto the

EARLY CRETACEOUS ORGANIC FACIES

adjacent slope and outer platform (Uliana & Legaretta 1993, p. 408). The facies also varies laterally, with the proportion of 'black shales' increasing towards the basin centre (Leanza 1981, p. 72; Macellari 1988, p. 402), becoming the dominant lithology in the north near Malargiie (Sagasti & Poire 1998; Uliana et al 1999, p. 25). Although the Agrio Formation extends up into the earliest Barremian, most of its thickness is Hauterivian (Leanza 1981, p. 72); the maximum total duration of deposition was approximately 10 Ma, from 136 to 126 Ma (Aguirre-Urreta et al. 1999, p. 37; Sagasti & Ballent 2002, p. 726). The mudrock-dominated Agrio Formation is typically divided into three members by the occurrence of a late early Hauterivian sandstone, the Avile Member, which forms a distinct topographic feature. This sandstone is commonly 10-40 m thick at outcrop (Legaretta & Uliana 1991, p. 438; Aguirre-Urreta et al. 1999, p. 35) but may reach 82m (Leanza 1981, p. 72), and can be as much as 200 m thick in the subsurface (Kozlowski et al. 1998, p. 20). It is a predominantly aeolian and fluviatile non-marine unit that was deposited as an extensive lowstand wedge during a sea-level fall; some marine influence has recently been identified in its lowermost part (Veiga et al. 2002, pp. 1002 and 1007). Although this regressive event may have lasted for as little as 0.25 Ma, it evidently had a profound, if temporary, effect across much of the lowgradient Neuquen Basin (Legaretta & Uliana 1991, p. 433). The Avile Member is an important reservoir rock in north-central Neuquen and southern Mendoza (Urien & Zambrano 1994, p. 523; Kozlowski et al. 1998, p. 20), but poor characteristics and poor seals have permitted only non-commercial oil accumulations in the vicinity of Chos Malal (Cruz et al 1998). The Agrio mudrocks below and above the Avile Member are now referred to as the Pilmatue and Agua de la Mula members, respectively (Leanza & Hugo 2001). In the Chos Malal area the Pilmatue Member is approximately 620630m thick (Cruz et al. 1996, p. 49, 1998; Aguirre-Urreta et al. 1999, p. 35). The latest early-late Valanginian transgression that initiated deposition of the Agrio Formation is regarded as one of the main stratigraphic turning points in the history of the Neuquen Basin (Legaretta & Uliana 1991, p. 442). Following flooding of the basin, black-coloured organic-rich sediments are best developed in the lower part of the Pilmatue Member (Riccardi 1988, p. 14; Cruz et al. 1996, p. 45; Kozlowski et al. 1998, p. 20). An organic-rich interval also occurs at the base of the overlying Agua de la Mula Member

253

(Uliana & Legaretta 1993, p. 408; Uliana et al. 1999, p. 26). Aguirre-Urreta et al. (1999, pp. 35 and 37) observed that these latest early Hauterivian shales (here informally referred to as the 'Spitidiscus shale') are distinct massive, dark bluish shales which mark an important global and regional marine transgression; this transgression was associated with rapid deepening (Veiga et al. 2002, pp. 1004, 1014 and 1018) and a temporary return to deposition of basinal organicrich facies (Legaretta & Uliana 1991, p. 438). The history of the Agrio Formation thus records two major transgressions (associated with organic-rich sediments), and at least one major regression (associated with the Avile Member). The stratigraphic summary diagram of Veiga et al. (2002, p. 1002) shows two intervals of progradation within the Pilmatue Member; the first of these would explain the change to the less organic-rich middle third of the member (see the next section), while the second culminates in the Avile Member. Veiga et al. (2002, p. 1002) show three progradational phases in the Agua de la Mula Member; three sequences, comprising just transgressive and highstands systems tracts, are also identified by Spalletti et al. (2001). Leanza (1981, p. 85) describes the basinal facies of the Agrio Formation as commonly bituminous 'dark grey to dark brown and black, thinly bedded, soft calcareous shales and marls'. Legaretta & Uliana (1991, p. 435) interpret the common occurrence of very thin tabular bedding and lamination, and the scarcity and low diversity of epifaunal and infaunal organisms, as indicating anaerobic-dysaerobic biofacies; anoxic or euxinic conditions are also sometimes inferred (Legaretta & Uliana 1991, p. 438), especially for the transgressive maxima during the late Valanginian and early Hauterivian (Uliana & Legaretta 1993, p. 397). The most common macrofossils are rare and scattered impressions of ammonites and inoceramid bivalves (Uliana et al. 1999, p. 20). In the Pilmatue Member the highest frequency of fossils is often associated with the occasional limestone and sandstone storm beds (Aguirre-Urreta et al. 1999). Sagasti & Poire (1998) describe a 78 m-thick section of the Pilmatue Member near Malargiie, and find that 18% by thickness was non-bioturbated and a further 44% poorly bioturbated, with evidence of greater oxygenation and current activity in the upper part of their section (reflecting the mid-Pilmatue progradation event?). Sagasti & Ballent (2002) observe that the transgressive Valanginian part of the Pilmatue Member is often rich in pyrite, organic

254

R. V.TYSON ETAL

matter and radiolaria, and is characterized by a monotypic foraminiferal assemblage; these are all features that suggest dysoxic bottom conditions, but the presence of Thalassinoides indicates that fully oxic periods also occurred, implying redox cyclicity. Sagasti (2000) notes that Milankovitch-scale rhythmic interbedding of shales and marls is characteristic of the distal Agrio Formation in Mendoza; carbonaterich beds in the 100-140 m-thick Agua de la Mula Member have 68-90% CaCO3 and carbonate-poor marly beds 42-65% CaCO3. On the basis of field observations (lithology, sediment fabric and ichnofaunas) Spalletti et al (2001) report a strong prevalence of anoxic, suboxic and dysoxic shale and marl facies (71-89% by thickness) within the Agua de la Mula Member at Loma La Torre (Fig. 1), with about 40% of the section corresponding to laminated, strongly fissile 'bituminous black shales' (Spalletti et al. 2001, p. 613), especially in the second and fourth quarters of the member (their sections B and E2), which are actually interpreted as more shallow-water highstand intervals.

Previous organic facies work As is commonly the case, the combination of dysoxic-anoxic conditions and sediment starvation (i.e. good preservation of marine organic matter and low siliciclastic dilution) favoured the widespread deposition of organic-rich oilprone source-rock facies during the transgressive intervals of the Agrio Formation (Legaretta & Uliana 1991; Uliana & Legaretta 1993; Uliana et al. 1999). Uliana & Legaretta (1993, p. 408) note that the total organic carbon (TOC) content of the Agrio black shale facies can be locally as high as 3-5%, although Urien & Zambrano (1994, p. 529) report a range of only 0.2-1.9% TOC (increasing northwards, p. 528). Uliana et al. (1999, p. 25) cite an intermediate mean TOC of 2-3% and mainly marine amorphous kerogen; they also present modified van Krevelen and S2 v. TOC diagrams for the Lower (Pilmatue) and Upper (Agua de la Mula) members of the Agrio Formation, with nonspecific data envelopes indicating hydrogen indices of 380-680 and TOC values of 1.58.0% in the former, and a somewhat lower quality of 320-580 and 0.9-3.5% in the latter (Uliana et al 1999, pp. 21 and 23). The most detailed published observations are those of Cruz et al. (1996, 1998) and Kozlowski et al. (1998), who describe the organic facies in three exploration wells drilled in 1995-1996 (including Chapua Este x-1, 16 km to the North of Chos Malal, and Pehuenche 165 km to the

North, Fig. 1), but based on only 18-24 samples from each. Cruz et al. (1996, p. 49) describe two organic facies within the 625-m thick Pilmatue Member, and Kozlowski et al. (1998, p. 20) note that this pattern is expressed regionally. The lower 200 m are characterized by almost exclusively marine amorphous kerogen, TOC values of 2.2-3.5% and hydrogen indices of 150-200, while the remainder is characterized by more mixed kerogen (up to 15% terrestrial contribution), lower TOC values of 0.8-1.8% and slightly higher hydrogen indices of 200-250. The shales of the lower part of the Agua de la Mula Member (described as dark grey rather than black) also apparently correspond to the latter of these two organic facies, but with HI values of 300-400 (two samples only). The difference in the two Agrio organic facies is masked by the thermal maturity gradient, which explains the divergence in the TOC and HI trends (HI being more affected by maturation than TOC); allowing for the maturity differences, the lower organic facies would definitely have been of originally higher quality. Vitrinite reflectance (VRo), spore colour and rmax values indicate the Agrio Formation has reached the oil window (probably in the Palaeogene), exhibiting a progressive increase in maturity with depth through the Chapua Este well, from early mature in the upper member (VRo <0.7%) to late mature (VRo 1.0-1.3%) at the base of the formation (a TmaK range of 435455 °C). This maturity range is a little lower than suggested by the vitrinite reflectance contour map given in Urien & Zambrano (1994, p. 528) (Fig. 1), even though the maturity in the Chapua Este well is considered to be locally elevated, perhaps due to the influence of Tertiary and Quaternary volcanism (Urien & Zambrano 1994, p. 532). Cruz et al. (1998) consider the lower and upper parts of the Pelmatue Member to have a higher source-rock potential; this is not particularly well demonstrated in the limited geochemical data for the Chapua Este well, but seems more apparent in the two more northerly wells. In all three wells a tripartite division certainly seems to be indicated by published gamma-ray logs (e.g. Cruz et al. 1996, p. 53); the hottest relative gamma-ray response actually occurs in the more calcareous lower third of the Agua de la Mula Member (including the 'Spitidiscus shale' interval). According to the data in Kozlowski et al. (1998, pp. 17 and 19), the subordinate terrestrial phytoclasts in the upper Pelmatue Member and lower Agua de la Mula Member probably become increasingly opaque (inertinitic) in a northerly, distal direction (an often-observed pattern: Tyson

255

EARLY CRETACEOUS ORGANIC FACIES

1995). Pramparo & Volkheimer (1999) report an exclusively terrestrial, phytoclast-dominated kerogen assemblage from clay intraclasts within the Avile Member at Cerro de la Parva (approximately 15 km NW of Chos Malal).

Sections studied The organic facies results reported here are based on 62 samples from five outcrop sections in the vicinity of Chos Malal (Table 1; Fig.l), from south (proximal) to north (distal) these are: (1) Pichi Mula (PM, contiguous with the Agua de la Mula section of Aguirre-Urreta et al 1999); (2) Lonco Vaca (LV); (3) Mina San Eduardo (MSB, see Aguirre-Urreta & Rawson 1993, p. 55); (4) Puerta Curaco (PC, see Leanza 1981, p. 64); and (5) Estancia Pampa Tril (EPT, on the east flank of the Las Yeseras anticline and on the west side of national road 40, north of the Loma la Torre section of Spalletti et al., 2001). Field reconnaissance and sampling was performed by the first author during 11-14 April 1999, with the stratigraphic guidance from M.B. Aguirre-Urreta (University of Buenos Aires) and P.P. Rawson (University College London). Sampling was focused mainly on the more organic-rich intervals at the base of the Pilmatue and Agua de la Mula members. Lithological sections were measured at PM, PC and EPT (the latter two in the company of E.G. Ottone of the University of Buenos Aires), but only reconnaissance sampling (of the darkest coloured intervals observed) was performed at LV because of the poor and weathered exposure, and the evident rarity of dark-coloured sediments. The 'Spitidiscus shale' (basal Agua de la Mula Member) sampled at PM was from a near-vertical face in a small gully developed above its sharp contact with the Avile Sandstone;

the ammonite Spitidiscus is also present immediately above the measured section where the shales apparently become softer and paler. Samples collected during logging of this section were supplemented by subsampling of samples through the Agua de la Mula Member previously collected by M.B. Aguirre-Urreta and E.G. Ottone, based upon the first author's examination of 70 palynostratigraphic slides prepared by Dr Ottone. Graphic logs are not shown because of the relative uniformity of the facies and the lack of meaningful stratigraphic variation in our data evident at this sampling density (but see those of Spalleti et al. 2001 from Loma La Torre). Samples of the Pilmatue Member at MSE and Agua de la Mula (AdM) come from the sets collected previously by M.B. Aguirre-Urreta and E.G. Ottone. Subsampling was, again, based on examination of palynostratigraphic slides prepared for 16 samples by Dr Ottone (only the three with conspicuous amounts of AOM, 5571m above the base of the Agrio, were selected). The subsamples selected from MSE comprise the darkest two field samples from the Pilmatue Member, plus all four samples taken from the 'Spitidiscus shale'. Although few samples were obtained from the Pilmatue Member at LV and MSE, the deliberate choice of the darkest layers (or samples) means that the data should reflect the highest quality organic facies developed at these localities; these atypical samples thus provide a conservative estimate of any decrease in organic facies quality towards the southern end of the sample transect.

Methods The samples were first cleaned to remove any contamination, and their lithology, fossil content

Table 1. Sections and lithostratigraphic units studied (from north to south, distal to proximal), and the number of samples from each. Sections followed by '(RVT)' are ones studied and sampled in the field by the first author. The remainder are sections studied and sampled by M.B. Aguirre-Urreta and E.G. Ottone (University of Buenos Aires) and subsampled as described in the text. Plus symbol after thickness indicates that the thickness is that examined, rather than the total thickness of the unit. Ss denotes the basal Agua de la Mula Member (the informal 'Spitidiscus shale') Code

Section

EPT PC MSE LV AdM PM PM

Estancia Pampa Tril (RVT) Puerta Curaco (RVT) Mina San Eduardo Lonco Vaca (RVT) Agua de la Mula Pichi Mula (RVT) Pichi Mula

Unit Pilmatue Pilmatue Pilmatue & Ss Pilmatue Pilmatue

Ss Agua de la Mula & Ss

Thickness

Samples

53.5 m + 65.3 m + 32.5 m (Ss) 3m + 242m 13m + 554m

12 9 2&4 3 4 12 20 & 2

R. V.TYSON ETAL

256

and their rock colour lightness values recorded. Rock-Eval-type pyrolysis was performed on 50-100 mg of powdered sample using a LECO THA-200 Thermolytic Hydrocarbon Analyser; SI, S2 and Tmax values were obtained (see Peters 1986). Total carbon and total organic carbon (TOC) analyses were performed using a LECO HE-100 Induction Furnace coupled with a LECO-244 Carbon/Sulphur Determinator; whole-rock carbonate values were obtained from the inorganic carbon content multiplied by 8.33, on the assumption that the carbonate was predominantly calcite. Following bulk geochemical analysis a subset of 22 samples was selected for transmitted light palynofacies analysis based on the observed range of hydrogen index values. Standard non-oxidative palynological preparation procedures were used (Barss & Williams 1973); the resulting residues were sieved at 20 jjim, strew mounted with Elvacite 2044 acrylic resin and examined using an Olympus BH2-RFCA microscope. Counting of 300 particles per sample was performed via manual traverses under a x 20 objective, recording the data on an electronic Swift Model F point counter. The count data represent the relative (%) numeric particle frequency of each kerogen category (Tyson 1995, p. 433); for this preliminary analysis, only a simple source-rock kerogen classification was used just to permit correlation between the optical kerogen character and the measured hydrogen indices (amorphous organic matter (AOM), brown and black phytoclasts, and undifferentiated palynomorphs). Where maturity permitted (mostly in samples from the Agua de la Mula Member), the fluorescence preservation scale of Tyson (1995) was applied, based on short blue-light irradiation; if no autofluorescence

was observed with a x 20 objective, the samples were examined with a x 40 objective to determine whether any residual fluorescence was present.

Results and discussion Bulk geochemistry and microscopy The mean geochemical characteristics of the Pilmatue Member, the 'Spitidiscus shale', and the overlying remainder of the Agua de la Mula Member are presented in the Appendix and summarized in Table 2. The mean organic carbon values of the lower two units are in the range 2-3%, similar to those reported previously for the Agrio Formation. The post-Spitidiscus Agua de la Mula Member (AdM), and also the Pilmatue Member samples from AdM, LV and MSE, are organic-poor (the mean TOC for both is 0.43%); these sections also exhibit lighter sediment colours (mean rock lightness values 3.44.6). The TOC is significantly influenced by the carbonate content, which is generally highest in the calcareous shales and marls of the 'Spitidiscus shale' and Pilmatue Member at PC and EPT (Fig. 2), and within any one section the correlation is generally negative (due to dilution). Samples with less than 30% carbonate show a mutual dilution of both carbonate and organic matter by clay, producing a positive correlation between carbonate and TOC (cf. Ricken 1993). The Agua de la Mula Member evidently has much lower carbonate values than have been reported for more distal and thinner sections in Mendoza (Sagasti 2000), suggesting higher siliciclastic dilution. The measured hydrogen indices suggest that the more organic-rich lower two units currently

Table 2. Summary of geochemical characteristics for samples from each of the stratigraphic units studied (the 'Spitidiscus shale' and the overlying remainder of the Agua de la Mula Member are treated separately). Unit

No. Lightness TOC Reactive Max. C03 TOC

Agua de la Mula Member Spitidiscus shale Pilmatue Member

14

4.6

0.43

41%

0.73

18 30

3.4 3.2

2.07 2.68

44% 67%

3.90 7.43

T Max. HI * max HI (slope)

S2

HI

AOM (n)

8

0.06

16

28

44

439

23(3)

51 27

3.57 4.72

195 141

541 285

387 174

449 455

87(9) 89 (10)

All values except Max. TOC and Max. HI are mean values; No. is the number of samples analysed; 'Lightness' is the measure of rock darkness from the Munsell colour scale (3 denotes dark grey and 5 medium grey); TOC is the total organic carbon (wt%); 'Reactive' refers to the proportion of TOC that is associated with generation of hydrocarbons upon pyrolysis (see text for explanation); CO3 is the carbonate (wt%) content derived from the inorganic carbon and assuming all carbonate is calcite; S2 is the yield of hydrocarbons (mg HC g"1 rock) generated by kerogen cracking during pyrolysis; HI is the mean of the individual observed hydrogen indices (S2 normalized to the sample TOC); 'HI (slope)' is the mean HI calculated from the slope of the regression line on a S2 v. TOC plot (which corrects for inert carbon and the carbon adsorbed and retained by the mineral matrix in organic-poor samples); rmax is the temperature of maximum generation of S2 hydrocarbons; AOM is the optically determined relative 'particle' frequency of amorphous organic matter in the samples, with '«' the number of samples studied

EARLY CRETACEOUS ORGANIC FACIES

257

Fig. 2. The relationship between total organic carbon and carbonate contents in the Agrio Formation. Note the presence of two trends: a positive correlation at less than 30% carbonate, suggesting mutual dilution by clay; and a negative correlation indicating dilution of clay and organic matter when carbonate exceeds 30%. The abbreviations for the rock units are P, Pilamtue; A, Agua de la Mula; Ss, 'Spitidiscus shale'. The locations are AdM, Agua de la Mula; PM, Pichi Mula; MSB, Mina San Eduardo; LV, Lonco Vaca; PC, Puerta Curaco; EPT, Estancia Pampa Tril.

have Type III kerogen. The kerogen assemblages for the Pilmatue Member and the 'Spitidiscus shale' are both strongly dominated by AOM, and, considering the level of maturation, this suggests their original genetic kerogen type was either Type II or II/III (organic facies B or BC sensu Jones 1987). The first author's quick visual examination of Dr Ottone's 70 slides from the whole Agua de la Mula Member at PM indicates that phytoclasts are dominant throughout, except within the 'Spitidiscus shale' during the the post-Avile flooding event. For those samples studied optically by us, a dominance of phytoclasts over AOM is associated with hydrogen indices of less than 100 and TOC values predominantly less than 1%. The overall mean ratio of black (opaque) to brown (translucent) phytoclasts is 2.0, indicating that the phytoclast material is mostly inert or semi-inert. The fraction of reactive or effective TOC (i.e. that associated with hydrocarbon generating organic matter) was approximated from the positive TOC intercepts on S2 v. TOC diagrams (i.e. means of [1-[intercept TOC/sample TOC]] x 100). The mean reactive fraction of the TOC is distinctly lower in the 'Spitidiscus shale' (44% v. 53-81% in the Pilmatue Member). Although the mean TOC for the

'Spitidiscus shale' is higher, the mean reactive TOC is actually lower (1.1% v. 2.3-2.8% for the Pilmatue Member at PC and EPT). Figure 3 shows a correlation between the logtransformed values of reactive TOC (on a carbonate-free basis) and the AOM : phytoclast ratio (r2 = 0.83, n = 13). For the organic-rich Pilmature member at EPT and 'Spitidiscus shale' the mean HI calculated from the slope of S2 v. TOC is around 300 or greater, indicating Type II kerogen (Fig. 4). The mean HI for the phytoclast-rich Agua de la Mula Member and the organic-poor Pilmatue Member samples (AdM, LV, MSB) is in the range 44-55, which equates to inertinitic, Type IV kerogen. The organic facies variation largely reflects a simple two end-member mixing between originally oil-prone AOM (Type II) and inert terrestrial phytoclasts (Type IV); as all the studied sections are relatively distal overall with respect to terrestrial inputs and palaeocoastlines, this mixing will largely reflect redox-controlled preservation of the AOM. Influence of maturity Some differences in the organic facies characteristics of the Pilmatue Member and 'Spitidiscus shale' can be attributed to their respective

258

R. V.TYSON ETAL

Fig. 3. The correlation between log-transformed values of the AOM: phytoclast ratio and the reactive TOC (wt%) expressed on a carbonate-free basis. Data from all units, but one atypical sample removed; reactive TOC could not be calculated for some samples where the observed TOC was less than the carbon intercept value on S2 v. TOC plots. For the key to the abbreviations see the footnote to Figure 2.

levels of organic maturation. The mean Tmax values of the Pilmatue Member correspond to approximately the middle of the oil window, while those for the 'Spitidiscus shale' indicate the early oil window (cf. Peters 1986). Because of low TOC values, only two Tmax values from the post-Spitidiscus Agua de la Mula Member were considered reliable, but the mean for these suggests that this unit is just within the oil window. Homohopane and sterane biomarker maturity data (not presented) also indicate that the Agua de la Mula Member is early mature. Visual fluorescence observations confirm a significant difference in maturity between the Pilmatue Member and the 'Spitidiscus shale' in the former the AOM was usually non-fluorescent (usually associated with a vitrinite reflectance greater than 0.7-0.8% according to Collins 1990, p. 42), whereas in the latter unit (at PM) the AOM was often weakly to distinctly fluorescent (yellow-greenish yellow colours). The higher maturity of the Pilmatue Member explains its lower pyrolysis S2 and hydrogen index values relative to the 'Spitidiscus shale' (despite the comparable TOC values and strongly AOM-dominated kerogen in both), and agrees with previous findings of Cruz et al. (1996, 1998) and Kozlowski et al (1998) based

on the adjacent Chapua Este well (Fig. 1; see above). Lateral variation The average values for geochemical parameters in the Pilmatue Member do not take into account the significant lateral variations in the characteristics of this unit. This variability is well demonstrated in Figure 4, which shows the S2 v. TOC plots for the Pilmatue Member differentiated by section, and the two intervals of the Agua de la Mula Member. The samples from the two most distal localities (EPT and PC) exhibit significantly higher slopes, with the mean HI calculated from the slope (Langford & Blanc-Valleron 1991) reaching a value of 297 at EPT, much closer to the 387 of the less mature 'Spitidiscus shale'. To better illustrate the lateral variation in geochemical characteristics of the Pilmatue Member, box plots were produced showing the median and quartile distributions for carbonate, TOC, carbonate-free TOC and hydrogen index (Fig. 5). There are relatively few data from the three most proximal sections (four for AdM, three for LV and two for MSE), but, as explained previously, the darkest available sediments or

EARLY CRETACEOUS ORGANIC FACIES

259

Fig. 4. The relationship between S2 and TOC in the Pilmatue Member (upper) and the Agua de la Mula Member (lower); data points are distinguished by lithostratigraphic unit. Note the different scales on the two plots. When multiplied by 100 the slope of the regression lines (the first term in the equations) gives the mean hydrogen index of the reactive or effective organic component

samples were selected for each to assess the maximum source potential. The carbonate content of the studied samples shows a progress!vestep-like increase from a median of about 5% in the two most proximal localities, to 25-30% in the next two and about 45% at EPT. This increase is also reflected in the median TOC values (not shown), which increase progressively from about 0.3% at AdM to about 3.5% at PC, but then decrease slightly to about 3% at EPT due to greater carbonate dilution. The median carbonate-free TOC values remove this effect and increase progressively in a distal direction, reaching about 6% at EPT. The hydrogen indices also increase, but less linearly, showing a much stronger increase in the two most distal sections studied. The latter two sections are also distinct in that the overall proportion of dark shale or marly layers is visually much

higher (being dominant at EPT and Lomo la Torre). Pyrolysis SI values (free hydrocarbons) also increase progressively in a distal direction, associated with a noticeably greater bituminous odour; the production index (S1/(S1+S2)) shows an opposite trend, with high values (>0.35) at AdM indicating traces of migrant hydrocarbons in otherwise organic-poor sediments. As there is no statistically clear difference in the maturity of the Pilmatue Member between the different localities (based on Tmax and other parameters), the observed variations are considered to reflect original depositional differences, and the transition from a largely dysoxic-anoxic basin or slope in the north (EPT and PC) and a predominantly oxic slope or outer platform environment in the south (AdM, LV and MSE). As the samples from the

260

R. V.TYSON ETAL

Fig. 5. Lateral variations in geochemical properties of the Pilmatue Member at the five studied localities, ranked from south to north (proximal to distal) succession: 1, Agua de la Mula; 2, Lonco Vaca; 3, Mina San Eduardo; 4, Puerta Curaco; 5, Estancia Pampa Tril. For 1 the data refer only to samples containing significant marine AOM; and for 2 and 3 they refer only to the darkest lithologies observed at outcrop (see text). The box plots show the median and quartiles of the distribution of each parameter.

most southerly (most proximal, but still relatively distal) sections were the darkest and most organic-rich observed, the true gradient in organic facies is even more marked than that shown. It is possible that a lateral transition from oxic to dysoxic-anoxic facies also occurs within the Agua de la Mula Member. The samples studied here, and examination of Dr Ottone's samples from the post-Spitidiscus Agua de al Mula

Member at PM, indicate prevalently oxic conditions (lighter colours, low TOC, low HI and the dominance of phytoclasts). This is in distinct contrast to the apparent lithological nature of the Agua de la Mula Member at Loma La Torre (about 10km south of EPT) as described by Spalletti et al (2001), who consider the shales and marls there to be generally dysoxic-anoxic facies. In agreement with the latter, Cruz et al. (1996) note that two samples from the Agua de

EARLY CRETACEOUS ORGANIC FACIES

la Mula Member in the Chapua Este well (Fig. 1) have hydrogen indices of 300-400, which would, indeed, imply dysoxic rather than oxic conditions. Maps showing the extent of the organic-rich facies have been published previously for the early and late Valanginian and the early and late Hauterivian (Legaretta & Uliana 1991, pp. 437 and 439; Uliana & Legaretta 1993, p. 407; Uliana et al 1999, pp. 21 and 23); the latter maps suggest that the southern edge of the organic-rich facies was located a little south of Puerta Curaco (Fig. 1). The observations present here are in broad agreement with these maps, but it is clear that the relatively organic-rich latest early Hauterivian 'Spitidiscus shale' extends at least as far South as Agua de la Mula (Pichi Mula); the presence of a benthic fauna in some samples (notably the gastropod Protohemichenopus neuquensis here confirms dysoxia rather than anoxia. The curtailed thickness of these organic-rich sediments at Pichi Mula (<13m), compared with the c.l 00m at Chapua Este (inferred from the gamma-ray log of Cruz et al. 1996, p. 53), may reflect the shallower and more oxic setting (where progradation would also have had a greater and more immediate impact).

Global records of Valanginian Hauterivian dysoxia-anoxia Macellari (1988, p. 409) observes that the Cretaceous anoxic events recorded in southern South America, although related to transgressions, do not coincide with the widely recognized 'global anoxic events' for this period (cf. Jenkyns 1980). In order to establish the larger context of the Agrio Formation, we therefore briefly review the general occurrence of organic-rich sediments in the Valanginian and Hauterivian. It should be noted that the relative rarity of black shales at this time is a reflection of the relatively low (but generally rising) global sea levels; the North Atlantic was relatively narrow, the South Atlantic was not yet open, the Western Interior Seaway of the USA was not yet established, and much of European shelf was characterized by non-marine Wealden or shallow carbonate-platform facies. Most shelf areas were evidently too shallow for stable watermass stratification and anoxia. Dark-coloured transgressive and partly dysoxic mudstone facies are developed in the marine Boreal shelf sea in the North Sea and northern Germany during the late Valanginian (Tyson

261

& Funnell 1987, p. 72; Mutterlose & Bornemann 2000, pp. 737 and 746), but in this region conditions were not favourable for true black shale deposition until at least the Barremian. However, further north, the basal part of the transgressive Pebble Shale on the North Slope of Alaska is apparently Hauterivian in age, and, locally at least, has Type II kerogen and a TOC content of about 4% (Keller et al. 2002). Although black shales of the same age seem less common elsewhere, the transgressive episodes recorded in the Agrio Formation (the diachronous early-late Valanginian base of the Pilmatue Member and the latest early Hauterivian base of the Agua de la Mula Member) certainly appear to correlate well with the major sea-level events recognized in better know sections in Europe (cf. Rawson 1994). The second and last of these events are known to be correlated with transgressive surfaces and important ammonite migrations in many parts of the world (Rawson 1993, 1994). South Atlantic, Austral Basin and Antarctic Zimmerman et al. (1987, p. 273) report that Valanginian oil-prone organic-rich shales occur at DSDP Site 511 on the Falkland Plateau and in the deeper water gulf off the South African margin; however, subsequent refinement of the dating suggests that the organic-rich facies at Site 511 can only be definitely dated as far back as the Barremian or latest Hauterivian (Mutterlose & Wise 1990; O'Connell 1990, p. 86). The Petroleum Agency of South Africa (2002) has, however, reported the presence of latest Valanginian and Hauterivian 'good quality source shales' in the Algoa and Gamtos basins on the South African margin. Berriasian-Hauterivian organic-rich sediments (including at least 21 m of Valanginian) have been recovered from ODP Site 692 on the East Antarctica slope offshore of Dronning Maud Land; these sediments are mostly laminated (occasionally bioturbated) and have a mean TOC of 8.6% (O'Connell 1990, pp. 73-75). They represent a distal dysoxic-anoxic facies, and have a mean hydrogen index of 432 and a kerogen assemblage with 70-85% AOM (Thompson & Dow 1990). Macellari (1988, p. 405) notes that partially anoxic dark greyblack laminated shales (Rio Mayer Formation and lateral equivalents) also occur in the Austral Basin of southern Argentina, associated with maximum flooding during the late Hauterivian (to Barremian). Uliana et al. (1999, p. 29) observe that, although these sediments contain

262

R. V. TYSON ETAL

only 0.5-2.0% TOC, the abundance of mainly marine organic matter and pyrite indicate reducing conditions. Together these observations point to a relatively common occurrence of Valanginian-Hauterivian dysoxic-anoxic shale facies within the deeper parts of the seaway that developed between southern Argentina, the Falklands Plateau, South Africa and Antarctica during the early opening of the South Atlantic, and also on the outwards-facing Pacific margin (ODP Site 692, see above). Organic-rich dysoxic Lower and mid-Valanginian interbeds with TOC contents of 2-3% have also been observed in the NW Pacific on the Shatsky Rise during ODP Leg 198 (Shipboard Scientific Party 2002).

Central Atlantic and European Tethys For the Central Atlantic, Cool (1982, p. 19) notes that a significant change occurs in the BlakeBahama Formation during the early Valanginian, after which mm-scale lamination becomes dominant. Summerhayes & Masran (1983, p. 475) also observe that 'sediments first became commonly laminated in the Valanginian ... as sea level was rising' (also Lini et al. 1992, p. 380). Wise et al. (1986) find that the late Valanginian sees the first occurrence of 'black to dark grey carbonaceous clay stone' interbeds at DSDP Site 603B in the NW Central Atlantic; these beds form up to 33% by thickness during the Valanginian and Hauterivian, but are interpreted as distal organic-rich turbidites, perhaps redeposited from an up-slope oxygen minimum zone (Wise et al. 1986, p. 41). Waples (1983, p. 969) considers that (initially scattered) definitive evidence of true deep-basin anoxia in the North Atlantic does not occur until at least the late Hauterivian; however, Summerhayes & Masran (1983, p. 476) note that on a carbonatefree basis, maximum TOC values actually increase progressively from the Berriasian to the Hauterivian (as sea level rises), reaching values of up to 6-10% (but mostly less than 2%). The first, sporadic Cretaceous occurrence of black shale interbeds in the European Tethys also occurs during the Valanginian. Lini et al. (1992, p. 375) observed that at this time the Maiolica Formation of the Southern Alps changes in character to grey limestone and marls with interbedded subordinate cm-thick black shales. A more detailed description of these sediments is given for the Lombardy Basin by Bersezio et al. (2002); they observe that some radiolarian-rich black shale beds, up

to 10cm thick, with 1-2% TOC and mainly amorphous kerogen, occur in the late Valanginian (to earliest Hauterivian). Their data indicate that the hydrogen indices are always low (<300, Type III), but, as is commonly the case, this reflects the partially oxidized condition of the AOM, rather than terrestrial organic matter. The first widely acknowledged bona fide dysoxic event in the European Cretaceous is the 'Faraoni Level', a thin (25-42 cm) but distinctive black shale and limestone unit that occurs in pelagic and hemipelagic facies within the lower part of the latest Hauterivian catuolli Subzone (angulicostata Zone, Chron CM4) of central and northern Italy, SE France and the Betic cordillera (Cecca et al. 1994; Baudin et al. 2002). In contrast to the Valanginian black shales, this unit has hydrogen indices that range from 240 to 580, indicating variable but better preservation of marine organic matter (Baudin et al. 2002, p. 7). The horizon is associated with only a minor 813C anomaly (Cecca et al. 1994, p. 566; Lini et al. 1992; Baudin et al. 1999, p. 488). An equivalent of the 'Faraoni Level' event has not yet been conclusively identified within the diamantensis Zone of the Neuquen Basin; the event may be present within the Agua de la Mula Member, but higher resolution sampling might be required to locate it, as it is believed to be only 100 ka in duration. The European time equivalent of the Neuquen 'Spitidiscus shale' is the sayni Zone (AguirreUrreta pers. comm. 2003), but this does not seem to be as distinct, although it is associated with the important mid-Hauterivian biogeographic event (Rawson 1993, 1994). Gulf of Mexico Apart from ODP Site 692, the best record of Valanginian deep-water sediments rich in marine organic matter occurs in Units III and IV from DSDP Site 535 in the eastern Gulf of Mexico (Buffler et al. 1984). The sediments at this site show a strong Milankovitch-scale cyclic alternation between dark calcareous shales and marls, and lighter grey laminated-bioturbated marls and limestones (Tyson pers. obs. 1981), with carbonate contents predominantly greater than 60%. The mean hydrogen index for the early Hauterivian-mid-Valanginian (470613m), computed from the slope of S2 v. TOC, is approximately 580, indicating well-preserved and immature Type II kerogen (raw data

EARLY CRETACEOUS ORGANIC FACIES

compiled from Buffler et al. 1984). The kerogen assemblage is dominated by AOM, which is generally 60-95% and averages more than 80% in dark grey beds with lightness values of 2 (Tyson pers. obs.). Carbonate-free TOC values increase progressively, but erratically, through the Berriasian-Valanginian to the mid-Hauterivian, from about 3 to 10%. Sample hydrogen indices reveal two peaks, one in the late Valanginian (c. 600 m) and the other in the early Hauterivian (c. 500 m), but TOC is more irregular due to variable carbonate dilution. The greatest overall frequency of laminated sediments increases upwards and peaks in the early Hauterivian (Tyson pers. obs. 1981). These observations suggest that the Gulf of Mexico was significantly more confined than the Central Atlantic during the Berriasian-Hauterivian, and experienced a fluctuating, but lower, mean level of oxygenation. Isotopic events The wider significance of late Valanginian-early Hauterivian black shales is emphasized by their temporal association with the first significant 613C isotopic anomaly of the Cretaceous (Lini et al. 1992; Weissert et al. 1998; Wortmann & Weissert 2000; van de Schootbrugge et al. 2000; Erba et al 2004). The S13C carbonate excursion has the form of a relatively sharp positive shift of about 3%c, followed by a gradual return to background values of about l%o; the organic carbon 613C curve is similar in form, but lags slightly behind the carbonate one (Lini et al 1992, p. 337; Wortmann & Weissert 2000). The overall duration of the excursion was about 5 Ma (Weissert et al 1998, p. 191), occurring mostly between the Tethyan verruosum (earliest late Valanginian) and the radiatus (earliest Hauterivian) zones, CM11-CM9 chrons (Weissert et al. 1998, p. 193), with the positive shift occurring during the late Valanginian CM11-CM12 chrons (Erba et al 2004). Note that the onset coincides with the midValanginian biogeographic event of Rawson (1993, 1994). The isotopic excursion has been linked to the combined effects of the Valanginian and Hauterivian transgressions: accelerated nutrient cycling (and consequent eutrophication) associated with 'greenhouse' climate conditions, the occurrence of black shales and, in particular, the widespread drowning of carbonate platforms (Weissert et al 1998). Late Valanginian platform drowning is also a feature of the Neuquen Basin (Mitchum & Uliana 1985, p. 271). To the present authors' knowledge, this isotopic event has not yet been documented in

263

the Neuquen Basin, but the outcrops and improving biostratigraphic control now available for the Agrio Formation make this an excellent area for such an undertaking. Conclusions The marine shales and marls of the Valanginian Hauterivian Agrio Formation outcropping in the vicinity of Chos Malal include two main organicrich intervals: one at the base of the Pilmatue Member (Valanginian), and the other at the base of the Agua de la Mula Member (late early Hauterivian). Significant lateral variation occurs in the Pilmatue Member, with dark organic-rich intervals being rare in the south but dominant at the northern (more distal) end of the studied transect; this trend is matched by a progressive increase in mean and maximum organic content and petroleum potential. Valanginian-Hauterivian black-shale facies are rare on a global basis; compared to other contemporary shelf seas, the semi-enclosed nature of the Neuquen Basin, combined with its subsidence regime, apparently made its central and northern parts more susceptible to the development of stable water-mass stratification and dysoxia-anoxia during transgressive episodes. While there is sedimentological and isotopic evidence for decreasing deep-water oxygenation and increasing carbon burial during the Valanginian and Hauterivian, the localized development of true black shales also appears to be mostly related to topographic isolation (e.g. South Atlantic region, Gulf of Mexico). R.V. Tyson thanks the British Council for funding field work in 1999 (via awards to M.B Aguirre-Urreta and P.P. Rawson). B. Aguirre-Urreta and G. Ottone are thanked for their help and hospitality. R.V. Tyson would also like to thank H. Villar, G. Sagasti and F. Baudin for kindly supplying reprints of their relevant work. The analytical work by P. Esherwood and K.A. Pattison was undertaken while the latter were in receipt of NERC MSc studentships. Professors Weissert and Bersezio are thanked for their reviews.

Appendix Data by section listed from north to south (distal to proximal). For parameters see the heading for Table 2. Height is the distance above the base of the section (base of the Agrio, or top of the A vile Member) where known. Some unreliable rmax values were discounted when calculating means.

264

R. V.TYSON ETAL

Section

Unit

Sample

Height (m)

Lightness

%CaCO3

%TOC

HI

1 T max

EPT EPT EPT EPT EPT EPT EPT EPT EPT EPT EPT EPT PC PC PC PC PC PC PC PC PC MSB MSB MSB MSB MSB MSB LV LV LV PM PM PM PM PM PM PM PM PM PM PM PM AdM AdM AdM AdM AdM AdM AdM AdM AdM AdM AdM AdM AdM AdM AdM AdM AdM AdM AdM AdM

Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Spitidiscus shale Spitidiscus shale Spitidiscus shale Spitidiscus shale Pilmatue Pilmatue Pilmatue Pilmatue Pilmatue Spitidiscus shale Spitidiscus shale Spitidiscus shale Spitidiscus shale Spitidiscus shale Spitidiscus shale Spitidiscus shale Spitidiscus shale Spitidiscus shale Spitidiscus shale Spitidiscus shale Spitidiscus shale Spitidiscus shale Spitidiscus shale Agua de la Mula Agua de la Mula Agua de la Mula Agua de la Mula Agua de la Mula Agua de la Mula Agua de la Mula Agua de la Mula Agua de la Mula Agua de la Mula Agua de la Mula Agua de la Mula Agua de la Mula Agua de la Mula Pilmatue Pilmatue Pilmatue Pilmatue

EP-12 EP-11 EP-10 EP-9 EP-8 EP-7 EP-6 EP-5 EP-4 EP-3 EP-2 EP-1 LAG9 LAGS LAG7 LAGS LAG6 LAG4 LAGS LAG2 LAG1 SE105 SE113 SE115 SE116 SE4-1 SE5-1 LV1 LV2 LV3 S-09 S-10 S-ll S-08 S-07 S-06 S-05 S-04 S-03 S-02 S-01 S-01A MU-111 MU-105 MU-170 MU-166 MU-165 MU-157 MU-155 MU-141 MU-136 MU-133 MU-130 MU-128 MU-127 MU-125 MU-120 MU-115 MU7-2 1175 MU7-2 1176 MU8-1 1177 MU9-2 1186

47.3 38.5 32.1 24.8 14.5 11.0 9.0 6.3 4.7 3.5 1.7 0.8 64.3 56.3 38.8 32.3 22.3 10.3 9.1 4.1 0.9

4 4 3 3 4 4 3 3 3 3 2 2 3 4 3

49.6 12.6 14.8 31.1 18.7 38.3 44.9 66.4 59.0 52.1 45.5 61.7 25.0 33.3 25.0 0.0 0.0 16.7 25.0 25.0 16.7 25.0 25.0 33.3 25.0 25.0 25.0 25.0 25.0 25.0 46.9 44.8 61.8 55.1 65.4 53.6 51.4 58.2 46.0 20.1 30.0 59.6 43.3 25.0 25.0 20.4 25.0 9.7 41.7 0.9 33.3 33.3 33.3 33.3 7.0 33.3 33.3 41.7 33.3 33.3 33.3 33.3

2.76 4.94 1.91 5.18 4.49 2.85 2.51 2.93 2.87 3.20 4.10 2.45 2.09 1.11 3.83 4.48 7.43 4.28 1.97 1.64 3.65 3.87 2.70 2.83 1.95 2.30 2.48 1.09 1.51 1.25 3.13 3.90 1.10 2.47 1.07 0.71 2.68 2.64 1.63 2.06 0.57 0.58 1.93 1.49 0.49 0.27 0.47 0.49 0.24 0.41 0.40 0.51 0.24 0.32 0.46 0.56 0.73 0.42 0.32 0.23 0.20 0.30

233 212 189 283 285 201 236 259 239 236 280 250 135 77 139 141 169 89 114 107 78 134 201 205 149 68 44 42 22 19 269 40 492 260 541 118 240 135 83 4 76 134 80 149 5 13 6 10 12 14 9 18 10 9 8 28 24 17 22 17 20 13

461 453 451 448 460 449 452 453 460 456 453 451 453 449 453 456 465 456 455 463 456 446 440 431 434 458 443 466 447 451 428 443 443 438 448 443 443 443 494 496 443 430 504 435 508 516 503 531 471 519 500 469 501 486 524 436 444 465 455 507 479 498

5.3 4.4 3.1 2.1 1.3 1.2 1.1 1.0 0.5 0.4 0.3 0.1

534.0 486.8 483.8 446.5 423.0 303.0 253.8 227.5 189.0 139.3 132.3 115.3 64.0 29.5

2 3 3 2 3 3 4 3 3 3 3 3 3 3 4 3 3 3 3 3 3 4 4 4 4 4 3 3 6 3 6 5 6 4 4 4 4 6 4 4 5 4 4 4 4

%AOM

92 97

99 99

97 89 99 98 99 98 75 96

94

49 80 94 83 93

5 18 45 44

EARLY CRETACEOUS ORGANIC FACIES

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L.G. & BURNS, S.J. 2000. Paleoceanographic changes during the early Cretaceous (Valanginian-Hauterivian): evidence from oxygen and carbon stable isotopes. Earth and Planetary Science Letters, 181, 15-31. VEIGA, G.D., SPALLETTI, L.A. & FLINT, S. 2002. Aeolian/fluvial interactions and high-resolution sequence stratigraphy of a non-marine lowstand wedge: the Avile Member of the Agrio Formation (Lower Cretaceous), central Neuquen Basin, Argentina. Sedimentology, 49, 1001-1019. WAPLES, D.E. 1983. Reappraisal of anoxia and organi richness, with emphasis on Cretaceous of North Atlantic. AAPG Bulletin, 67, 963-978. WEISSERT, H., LINI, A., FOLLMI, K.B. & KUHN, O. 1998. Correlation of early Cretaceous carbon isotope stratigraphy and platform drowning events: a possible link? Palaeogeography, Palaeoclimatology, Palaeoecology, 137, 189-203. WISE, S.R., JR., VAN HINTE, I.E. ETAL. 1986. Mesozoic Cenozoic clastic depositional environments revealed by DSDP Leg 93 drilling on the continental rise off the eastern United States. In: SUMMERHAYES, C.P. & SHACKLETON, NJ. (eds) North Atlantic Palaeoceanography. Geological Society, London, Special Publications, 21, 35-66. WORTMANN, U.G. & WEISSERT, H. 2000. Tying platform drowning to perturbations of the global carbon cycle with a 613Corg-curve from the Valanginian of DSDP Site 416. Terra Nova, 12, 289-294. ZIMMERMAN, H.B., BOERSMA, A. & McCov, F.W. 1987. Carbonaceous sediments and palaeoenvironment of the Cretaceous South Atlantic Ocean. In: BROOKS, J. & FLEET, A.J. (eds) Marine Petroleu Source Rocks. Geological Society, London, Special Publications, 26, 271-286.

Palaeoclimatic implications of Middle Jurassic (Bajocian) coniferous wood from the Neuquen Basin, west-central Argentina HELEN S. MORGANS-BELL1 & DUNCAN McILROY2 1 Department of Earth Sciences, Oxford University, Parks Road, Oxford OX1 3PR, UK 2 Department of Earth Sciences, Memorial University of Newfoundland, St John's, NL, Canada A1B 3X5. Abstract: Silicified coniferous wood is commonly found in the Middle Jurassic (Bajocian) Lajas Formation of the Neuquen Basin, west-central Argentina. The wood is preserved in a succession of sandstones, siltstones, mudstones and minor conglomerates that represent deposition as part of tide-dominated deltas and fluvial plains across which large rivers meandered. Most of the wood occurs as dense accumulations in the tidal and fluvial channels. The wood fragments are worn, abraded, and lack both bark and branches, indicating that they were transported prior to deposition. The material is typically 20-30 cm long, with only infrequent examples of larger trunks (c. 80 cm in diameter, 5-6 m long). No trunks were found with root systems attached, and no stumps were found upright and in situ. The fossil wood genus Araucarioxylon dominates the assemblage. Growth rings are largely absent from the specimens, although one sample (from Rhea Gorge) displays highly diffuse and irregularly spaced rings, suggesting that it grew in different conditions from the others studied. Large-scale interpretations for southern Gondwana suggest a seasonally dry climate. However, these fossil wood specimens show no evidence of this, indicating that in this area at least the effects of any seasonal component to the climate may have been overridden by factors such as a locally plentiful supply of water and/or the possibility that growth was to some extent controlled by the taxonomic affinity of the trees.

The fossil wood-bearing Middle Jurassic (Bajocian) Lajas Formation of the Neuquen Basin in Patagonia, west-central Argentina (Fig. 1) comprises approximately 500 m of well-exposed tide- and fluvial-dominated facies that form unconformity bounded sequences (e.g. Gulisano & Hinterwimmer 1986; Zavala 1996; Mcllroy et al 1999, 2005). The deposits are interpreted to have been laid down as a series of deltas that prograded into a broad, structurally defined embayment that was located on the southern margins of the back-arc Neuquen Basin (Ramos 1988; Legarreta & Uliana 1991, 1996, and references therein; Riccardi et al. 1992; Vergani et al. 1995). Previous studies have shown that during the Mid-Jurassic the basin was limited to the west by an active volcanic arc, and to the south and NE by a landmass composed of Palaeozoic-Triassic volcanics and plutonics (e.g. Gust et al 1985; Riccardi et al 1992; Legarreta & Uliana 1996). The delta complexes that form the basis to this study are interpreted to have prograded into the basin from landmasses to the south and SE (Uliana & Biddle 1987; Burgess

et al 2000; Mcllroy et al 2005). The basin lay approximately 50°S during the Middle Jurassic (Ziegler et al 1993), and tectonic subsidence of the basin took place intermittently during the Jurassic as a whole, interspersed with phases of uplift across distinct fault blocks and/or the region as a whole (Burgess et al 2000). The specimens of conifer wood that make up this study were collected during a field season to central and southern parts of the Neuquen Basin in 1999, which in part aimed to characterize flood-plain and delta-plain facies across the region (see Mcllroy et al 1999, 2005). The wood was gathered from the area of Sierra de Chacaico, 50 km south of Zapala in Neuquen Province (Fig. 1). Apart from the fossil wood collected at the Fortin Primero de Mayo locality, which is remote from and poorly correlated with the main section, the stratigraphic positions of the other wood specimens are marked on Figure 2. This study utilizes six specimens that were collected to give an initial impression of the Middle Jurassic flora and palaeoclimatic regime of the Neuquen Basin. The investigation

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 267-278. 0305-8719/05/$15.00 © The Geological Society of London 2005.

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Fig. 1. Map of the Neuquen Basin and vicinity, west-central Argentina showing geographical features and the position of the localities from which the wood samples for this pilot study were collected. These localities are numbered: 1, Arroyo Carreri; 2, Cactus Gorge; 3, Maquina Cura Norte; 4, Rhea Gorge; 5, Estancia Charahuilla; 6, Fortfn Primero de Mayo. The stratigraphic position of localities 1 -5 is shown in Figure 2, but locality 6 is not shown because Fortin Primero de Mayo is remote from the main section and stratigraphic correlation is uncertain. The smaller index map shows the location of the study area in relation to the Neuquen Basin as a whole, with major rivers and province boundaries marked.

presented herein involves taxonomic and growth-ring studies. While the sample size is small, the paucity of other palaeoclimatic data for this interval in Argentina makes the study significant. Current environmental and palaeoclimatic interpretations are essentially based on palynological studies (e.g. Martinez et al. 1996; Quattrocchio et al. \996b). These authors recognize three separate, successive palynological assemblages in strata of Bajocian age, which they interpret as representing a warm-coolwarm climatic fluctuation. The current study is also important in providing a palaeoclimatic context for the Lajas depositional system, which is important if it is to be used as an analogue to other tide-dominated depositional

systems (e.g. the Tilje and He formations of offshore Mid-Norway: Martinius et al. 2000; Mcllroy 2004). Other regions of Patagonia distant to the Neuquen Basin yield abundant fossil wood (including silicified stumps, many in growth position, logs, branches and twigs of various sizes, seedlings and seed cones) and numerous studies have been conducted on the remains since their initial scientific discovery by Windhausen in 1919 and subsequent description by Gothan (1925). Calder (1953) provides a full history of the early investigation into petrified forests of Patagonia, situated around the volcanic peaks of Cerro Cuadrado, Cerro Madre e Hija and Cerro Alto in Santa Cruz Province, approximately

Fig. 2. Composite log of the Lajas Formation showing the variety of facies comprising the succession, the relative thickness of the constituent parasequences and the stratigraphic position of sequence boundaries (SB). The stratigraphic levels from which fossil wood was collected are numbered as in Figure 1. Note that the exact horizon that the Fortin Primero de Mayo locality represents is unclear. Inset bottom right is the general stratigraphy for the region, showing the position of the Lajas Formation, which is Bajocian in the study area but may be as young as middle Bathonian in age to the north.

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900 km to the south of this Neuquen Basin study area (see also Spegazzini 1924; Windhausen 1924; Riggs 1926, p. 544; Calder 1953). The material is preserved in beds of rhyolitic ash understood to be of Middle Jurassic age (sensu Stipanicic & Bonetti 1970). Taxonomic work on the wood fragments and seed cones has revealed the predominance of araucarian-type wood (e.g. Riggs 1926, p. 544; Wieland 1935, p. 6; Calder 1953), with some remains displaying indeterminate rings (Calder 1953, p. 105) that may represent annual growth (Wehrfeld 1935, p. 120). There are reports of trunks up to 100 m in length at Cerro Cuadrado (Wehrfeld 1935, p. 125), including some trunks with recorded circumferences of over 10m (Leanza in Feruglio 1950, p. 129, footnote). Today, only exceptional specimens of the North American Redwood (Sequoia sempervirens) grow to a height exceeding 100 m (Brockman 1986). The wood taxa described in this study are compared with this rich dataset and provide an initial impression of the Mid-Jurassic flora of the Neuquen Basin.

Geological setting of the wood samples The six fragments of silicified wood on which this study is based were collected from discrete levels in the upper part of the Lajas Formation (Fig. 2). The unit is primarily composed of sandstones, siltstones and mudstones, with minor conglomerates and shell beds, and these sediments make up several facies associations that include prograding tidal deltas, stacked sandy tidal channel fills, extensive bayfill deposits and tidal-flat successions, which are locally cut by heterolithic tidal-channel-fill facies (Mcllroy et al. 2005). Within the succession, fossil wood mainly occurs as allochthonous accumulations in tidal or fluvial channels, which are particularly common immediately beneath the junction with the overlying fluvially dominated Challaco Formation (Fig. 2). Field observations note that the wood is worn and abraded, lacking bark and branches, and entrained within tidal and fluvial channels, indicating that the material was transported prior to deposition. This is supported by the absence of trunks with attached branches and root systems, and lack of in situ tree stumps. The larger branches and trunks are found aligned with sedimentary current indicators as a lag to fluvial channel bodies or as chaotically organized log-jams. The woody material is typically 20-30 cm long, with infrequent examples of larger trunks (c. 80 cm in diameter, 5-6 m long). The wood is thus interpreted as being transported to the basin from

the drainage basin and, although allochthonous, provides information of regional importance. The tidal channels are characterized by up to 5 mm-thick organic-rich clay drapes and a marine fauna represented by traces of echinoderms (Scolicia isp. and Asteriacites isp.), along with a worm-like infauna (e.g. Dactyloidites isp.) and wood-boring bivalves (Teredolites isp.). The great importance of tidal processes during deposition of the Lajas Formation is unusual, as they are usually only predominant in incised valleys during rising relative sea level (see Dalrymple et al. 1992; Zaitlin et al 1994). Mcllroy et al (1999, 2005) relate the dominance of tidal processes to the funnel shape of the Neuquen Basin structural embayment, which succeeded in amplifying tidal currents during all states of relative sea level. Incised valleys are static features passively filled during early sealevel rise and usually overtopped by highstand deposits (Van Wagoner et al 1988). In contrast, structural features such as the embayment in which the Lajas Formation was deposited enabled accommodation space to be created continuously during deposition of the formation, thus maintaining tidal amplification throughout complete cycles of sea-level change. Fluvial channels in the basin are distinguished on the basis of coarse grain sizes that reflect the predominance of sand and gravel-sized clasts, which comprise lateral-accretion bedsets and trough cross-bedded channel fills. They also typically lack bioturbation. In places within the fluvial channels, tidal bundles rich in organic material occur. In macrotidal settings with an extensive delta plain, the tidal reach may extend many tens of kilometres up the fluvial channel, depositing organic-rich drapes at the high-tide slack water due to slowing of fluvial currents by interaction with opposing tidal currents. For example, the tidal reach extends 250 km into the delta plain in the case of the Fly River, Papua New Guinea (Baker et al 1995). The Lajas Formation is diachronous across the region and so the age of the unit varies from place to place. Recent work on spores, pollen and dinoflagellate cyst assemblages from the formation and the units it is sandwiched between suggest an age extending from the early Bajocian to the early-late Callovian (e.g. Volkheimer & Musacchio 1981; Quattrocchio & Volkheimer 1990; Quattrocchio et al 19966; Zavala 1996). The underlying Cura Niyeu Formation contains ammonites and palynological data that demonstrate an early-late Bajocian age, whereas beds within the uppermost Lajas Formation include microflora and dinocysts indicative of the early-middle Callovian (e.g. Quattrocchio

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et al. 19960, b). Dinoflagellate cyst assemblages examined as part of this study provide further age constraints: the Burgess Parasequence of the Cura Niyeu Formation is given a latest early Bajocian age (humphriesianum Zone), and the Owl Parasequence in the upper part of the Lajas Formation an earliest Late Bajocian age (niortense Zone: J. Fenton pers. comm. 1999). These new dates suggest that the 500600 m-thick Cura Niyeu and Lajas formations were deposited entirely during the Bajocian, not necessarily ranging into the Bathonian and Callovian as previously asserted. The period of deposition of the Lajas Formation is approximately 4.5 Ma based on available biostratigraphic evidence, which gives a mean sedimentation rate of 122 m/Ma~ .

Materials and methods Three mutually perpendicular planes of section were prepared for each sample to demonstrate the wood anatomy (transverse, TS; radial longitudinal, RLS; tangential longitudinal, TLS).

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The thin sections are lodged in the Oxford University Museum of Natural History (JU.153JU.158). To most effectively examine the specimens, petrological thin sections were analysed microscopically using reflective light and bright-field illumination, combined with a monochromatic green filter. The degree of replacement by silica varies and in some of the samples, or parts of the samples, the silicification is so extensive that little anatomical detail remains (e.g. as in sample JU.157 shown in Fig. 3b). In other parts of the same specimen, however, the cell-wall structure is well defined and anatomical details such as ray height (Fig. 4 g, h, sample JU.157) and arrangement of pits on the tracheid walls can be recognized (e.g. Fig. 4a-c, sample JU.157). In the examination of growth rings in the transverse plane, areas of distortion and cracks were avoided where possible and deformed rings were traced laterally to relatively unaffected areas where measurements could be resumed. Particular attention was paid to the occurrence of hairline fractures running parallel to the

Fig. 3. Range of preservation states and character of ring boundaries. The scale bar is 0.1 mm. (a) JU.156/p3: faint growth-ring boundary (arrowed); the transition from latewood to early wood appears gradual and is highlighted by a change from thicker- to thinner-walled cells; Rhea Gorge, (b) JU.157/p3: transverse section showing how the growth of quartz crystals has disrupted the cellular structure; Estancia Charahuilla. (c) & (d) JU.158/p3: examples of distortion and shearing; Fortm Primero de Mayo.

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Fig. 4. Anatomical detail of the radial and longitudinal planes. (1-4) Arrangement of bordered pits on radial walls of tracheids; pits are primarily biseriate, contiguous, alternate and polygonal, suggesting an araucarian character, less commonly the pits are unseriate (see c). (a)-(c) JU.157: Estancia Charahuilla. (d) JU.155: Maquina Cura Norte. (e) & (f) JU.154: mix of bordered pits arrangement, primarily uniseriate, contiguous and elongate in the transverse plane; Cactus Gorge, (g) & (h) JU.157: uniseriate rays and variable in height, 3-25 cells; Estancia Charahuilla. The scale bar is 0.1 mm.

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growth rings, because they can obscure or eliminate parts of, even the entire, ring. The samples from Arroyo Carreri (JU.153) and Cactus Gorge (JU.154) display some ring curvature and might therefore be from branches or small diameter trunks. The other samples display outer rings that have relatively straight boundaries when viewed in the transverse plane, and may therefore be derived from larger branches or trunks. In all samples, the growth rings show little variation in width when traced laterally. Distorted and crumpled areas of cells are common features of this dataset, particularly in the early wood of the growth rings. Early wood is susceptible to compaction and crushing because of its large, thin-walled cells that have little strength compared to the smaller, thickerwalled late wood cells. The amount and extent of crushing depends not only on relative cellwall thickness, but also on the degree of microbial degradation, which can be related to waterlogging or desiccation (cf. Jefferson 1982; Francis 1984; Creber & Chaloner 1985; Parrish & Spicer 1988) (Fig. 4c, d). Waterlogging may well be the main cause of the cellular distortion given the burial of the Lajas wood samples within tidal and fluvial channels. However, the effects of desiccation cannot entirely be ruled out, particularly given the presence of cell-wall checking in the Estancia Charahuilla sample (cf. Cope 1993; Jones 1993). Once entombed within the channel sands, silicification of the secondary wood probably took place during early diagenesis. During diagenesis, the high proportion of siliceous volcanic ash in the sediment (typically 25-30%, as determined by petrographic analysis carried out as part of this study) would have released free silica, allowing silicification of the wood. Large petrified logs and branches found in the channel facies are primarily preserved in this manner, thereby preserving a variable amount of anatomical detail. Throughout the rest of the Lajas facies, smaller, fragmentary woody material is common and is typically coalified, completely lacking in cellular structure. Fossil wood taxonomy While this information is of limited use in determining the precise taxonomic affinity of the tree (cones and leaves are more diagnostic) it does allow comparison of similar woods. The study focuses on features preserved in the radial and tangential longitudinal planes (Fig. 4). The variable and commonly poor preservation of the wood that comprises this dataset precludes a thorough investigation, and no useful taxonomic

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information whatsoever could be retrieved from the Fortin Primero de Mayo specimen. Simple observations on those samples that are well enough preserved are reported here. Wood MorphogenussAraucarioxylon. Kraus, in Schimper 1870 Material. This taxon is mainly described from the Estancia Charahuilla (JU. 157) and Maquina Cura Norte (JU. 155) samples. The Arroyo Carreri (JU.153) and Rhea Gorge (JU.156) specimens share similar characteristics, but are relatively poorly preserved. All samples are from the Sierra de Chacaico. Description. Annual ring boundaries generally indistinct, with a gradual transition from the earlywood to latewood. Bordered pits on radial longitudinal walls of the tracheids are biseriate, alternately arranged, mutually compressed and hexagonal in outline (Fig. 4a, c, d, JU.157). A subordinate number of contiguous, evenly spaced, uniseriate pits are also present. Cross-field pits are poorly preserved and commonly distorted in the diagonal plane. Medullary rays are mainly uniseriate, occasionally with a few biseriate cells within the single row strings. There is a mix of short and tall rays that are two-five or 20-25 cells in length (Fig. 4 g, h, JU.157). Remarks. Wood specimens showing this anatomy make up 80% of the collection. The samples are assigned to Araucarioxylon because of the presence of diagnostic multiseriate bordered pits on the radial walls of the tracheids (Fig. 4b), which are alternately disposed and polygonal in shape. This araucarian style of pitting corresponds to the anatomy of Araucaria and some other members of the modern Araucariaceae (Seward 1919; Krausel 1949). Fossil wood and seed cones related to Araucaria have previously been identified to the south of the Neuquen Basin in Santa Cruz Province. For example, Calder (1953) distinguished seed cones from Cerro Cuadrado as Araucaria mirabilis (Spegazzini) Calder and Pararaucaria patagonica (Wieland). In the fossil wood sample from Estancia Charahuilla, both 'araucarian' and 'non-araucarian' (circular, rounded, spaced pits) pits occur.

Taxon 1 Material. Single sample (JU.154, Fig. 4e, f) from Cactus Gorge, Sierra de Chacaico.

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Description. Bordered pits up to 25 jjim in diameter are uniseriate and contiguous, and in places elongate in the horizontal plane. Crossfield pits are poorly preserved, but there appear to be one-two pits in some fields. Rays are uniseriate and variable in height, between three and 25 cells in length. Remarks. The lack of diagnostic araucarian bordered pits distinguishes this wood from the rest of the sample collection. It is possible that the apparent absence of araucarian pits is simply an artefact of the specimen's preservation. Alternatively, the closely spaced, uniseriate pits that are elongate in the horizontal plane might be indicative of another taxon, quite separate from Araucarioxylon. Indeed, wood samples showing a similar arrangement of bordered pits were noted by Calder (1953) in specimens from Cerro Cuadrado and identified simply as 'coniferous wood'. Growth-ring analysis In hand specimen, the woods studied show subparallel, concentric, dark-coloured bands that suggest the presence of growth rings. Under the microscope, however, these apparent rings in fact correspond to distinct bands of sheared and deformed cells that roughly conform to the circumference of the tree, crumpling large sections of cells and obviating the possibility of measuring ring widths accurately. The samples from Estancia Charahuilla (JU.157) and Fortin Primero de Mayo (JU.158) are largely preserved in this manner (Fig. 3c, d). In the specimens that are better preserved, subtle changes in cell-wall thickness can be studied for evidence of interruptions in seasonal growth. Samples from Estancia Charahuilla (JU.157), Maquina Cura Norte (JU.155) and Cactus Gorge (JU.154) have a more-or-less featureless appearance in transverse section indicating that the trees from which they were derived grew in conditions that varied little on an annual basis, and that growth may well have been continuous. The Estancia Charahuilla specimen shows a slight perturbation, however, in the form of a band of three-eight cells that appear smaller and thicker walled than the earlywood cells to either side, and may represent a false ring (occurs before the first crumple zone in the thin-section). False rings represent a slowing or suspension of growth and can be attributed to a variety of palaeoenvironmental conditions including drought, waterlogged conditions, freezing or insect attack (Fritts 1976). The Rhea Gorge (JU.156) sample is different from those described above in that it contains,

across several centimetres of section, highly diffuse and irregularly spaced 'ring boundaries'. Fourteen diffuse ring boundaries occur within the 4 cm-long transverse section. Each ring measures less than 5 mm in width, or more commonly about 2 mm. The rings comprise between 22 and 75 cells, or 45 on average. When traced laterally the width of the rings is fairly regular. Within the earlywood of several rings (i.e. rings 2, 4, 5, 8, 9 and 10), crumpled zones of cells occur, distorting the cellular structure. In such cases, approximate cell counts could still be made. Ring width is important to gauge because it is a measure of the tree's productivity, where the wider the ring the longer the growing season or faster the growth (Fritts 1976). The annual variability of tree growth can be measured by the mean sensitivity calculation (Fritts 1976; Creber 1977). By definition, the mean sensitivity of a tree may vary between 0 and 2, representing no annual variation and great annual variation, respectively. These are, however, extremes rarely encountered in nature. The figure of 0.3 is used to split populations (Fritts et al. 1965): whereas a value of less than 0.3 distinguishes 'complacent' trees having little response to climate change, or growing in favourable conditions with little interannual climate change, such as in tropical, non-seasonal regions, a value of more than 0.3 suggests 'sensitivity' to climatic variation (Creber 1977). The rings comprising the Rhea Gorge wood yield a mean sensitivity value of 0.46, indicating that the tree was growing in an environment that varied significantly on an annual basis. Factors such as water availability, temperature and light are the main determinants influencing plant growth and, hence, mean sensitivity results, although local effects such as position of growth, disturbance during growth and taxonomic identity of the tree can also be important (Fritts 1976). The transition between the earlywood and latewood is extremely gradual in the Rhea Gorge sample (Fig. 3a), preventing accurate delimitation of the respective zones by eye. A more precise estimate of latewood width can be made using a technique devised by Creber & Chaloner (1984), which is based upon the measurement of cell diameters across individual rings. The average cell diameter is approximately 60 jjim in this sample. The method was applied to four successive, well-preserved rings in the Rhea Gorge sample, which are typical for the specimen. Figure 5 shows the results of this investigation. The mean of each suite of cell diameters was first determined, and then the cumulative sum of their deviations from the

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Fig. 5. Plot of radial diameters of cells across four successive rings from the Rhea Gorge (JU.156/p3) wood sample (shown by dotted lines). These rings are representative for the wood specimen as a whole. The number of cells comprising each ring range between 35 and 58. A five-point running average of the raw data collected from the four rings smoothes the results (grey line). For each ring, the average cell diameter was calculated. In turn, the cumulative sum of deviations from the mean was determined for each set of cell diameters. These were then averaged to give a representative suite of data. These are illustrated by the bold curve. The boundary of the early wood (EW) and latewood (LW) is indicated, based on the downturn of the bold line towards zero. The thin black line with a negative slope is a line of regression for the averaged cumulative sums. See text for discussion.

mean calculated (illustrated by the bold curve in Fig. 5). The peak of the line essentially represents the boundary between the earlywood and latewood, and judging by these results it appears that relatively unfavourable conditions for growth began midway through the season, producing a wide latewood zone (Fig. 5). These results correspond to categories D and E of the growth-ring templates of Creber & Chaloner (1984). Such ring types, especially type E, are typical of trees growing in regions where the seasons are fairly uniform. They are also associated with woods that show weakly developed or absent rings, as well as woods of the Araucariaceae (Creber & Chaloner 1984). Discussion This pilot study of coniferous woods suggests that Araucarioxylon dominates the wood assemblage from the Lajas Formation, Neuquen Basin. Three of the samples (from Estancia Charahuilla, JU.157; Maquina Cura Norte, JU.155; Cactus Gorge, JU.154) do not show any evidence for interruptions in growth, suggesting that the parent tree was not limited on a seasonal basis. In contrast, the weakly developed and irregularly spaced rings comprising the Rhea Gorge wood fragment indicates a rather different situation, where tree growth probably was to some extent

limited from season to season, although the diffuse nature of the ring boundaries hints that growth may have slowed rather than ceased between the main growing seasons. On the whole, what evidence can be recovered from the dataset indicates that growing conditions were good, and possibly year-round given the findings from three of our sample set. The age of these fossil wood samples has been interrupted as earliest Late Bajocian, an interval Quattrocchio et al. (I996b) link to warm-hot and humid conditions on the basis of palynological assemblages. These wood data fit with this general picture of the local palaeoclimate. However, broader studies of southern Gondwana conclude that the climate was seasonally dry at this time (e.g. Volkheimer 1970; Rees et al. 2000). If this is correct, one might expect fossil wood from the Neuquen Basin to display growth rings. Of course, it is quite possible that the climate was indeed seasonal, but that plants in the basin were able to grow continuously throughout the year because local factors were more important to plant growth than general climatic conditions. For example, a locally plentiful water supply might have allowed sustained growth. Similarly, if deep tap roots extended beneath the trees and into the water table below, growth might have continued throughout the year regardless of seasonal dry spells.

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Although this field study did not recover trunks with root systems attached to the worn and abraded wood fragments, the presence of tap roots is not unusual in members of the Araucariaceae: they have been found attached to trunks of Araucarioxylon arizonicum in the Petrified Forest, Arizona (Ash & Creber 1992). Another possibility is that genetic make-up of the trees might have influenced their growth. This point is especially relevant here as studies of modern conifer families indicate that growth rings are least conspicuous in the Araucariaceae, irrespective of climatic conditions (e.g. LaMarche 1982, p. 4; Falcon-Lang 2000). Assuming that it is feasible to relate the character of extant Araucariaceae to extinct Araucarioxylon, then taxonomic affinity may well have played a role in tree development. In summary, the wood samples indicate favourable growing conditions. Although this pilot study did not bring to light any evidence to indicate a climate with marked seasons (cf. Volkheimer 1970; Rees et al 2000), that is not to say that such climatic conditions did not exist, as they might have been masked in these wood remains by local or taxonomic factors. Further work on a larger collection of fossil wood is required to resolve better the relative importance of seasonal rainfall, water supply and genetic factors on tree growth during the Middle Jurassic in the Neuquen Basin. The fossil wood samples were collected as part of a study funded by Statoil, the Heidrun License and Saga Petroleum, with logistical support from Repsol YPF. We thank E. Joven (University of Liverpool) who assisted with field work, and O. Green (University of Oxford) who provided technical support in the microscopic study of the wood samples. Palynological data are from an unpublished study carried out by Dr J. Fenton of Robertson Research International Ltd and commissioned by Saga Petroleum. Valuable critical comments were provided by G. Creber, H. Falcon-Lang, J. Francis and A. Zamuner, for which we are extremely grateful.

References ASH, S.R. & CREBER, G.T. 1992. Palaeoclimatic interpretation of the wood structures of the trees in the Chinle Formation (Upper Triassic), Petrified Forest National Park, Arizona, USA. Palaeogeography, Palaeoclimatology Palaeoecology, 96, 299-317. BAKER, E.K., HARRIS, P.T., KEENE, J.B. & SHORT, S.A. 1995. Patterns of sedimentation in the macrotidal Fly River delta, Papua New Guinea. In: FLEMMING, B.W. & BARTHOLOMAE, A. (eds) Tidal Signatures in Modern and Ancient Sediments. International Association of Sedimentologists, Special Publications, 24, 193-211.

BROCKMAN, C.F. 1986. Trees of North America: A Guide to the Major Native and Introduced Species North of Mexico. Golden Books, New York. BURGESS, P.M., FLINT, S. & JOHNSON, S. 2000. Sequence stratigraphic interpretation of turbiditic strata: An example from Jurassic strata of the Neuquen basin, Argentina. Geological Society of America Bulletin, 112, 1650-1666. CALDER, M.G. 1953. A coniferous petrified forest in Patagonia. Bulletin of the British Museum (Natural History), Geology, 2, 99-138. COPE, M.J. 1993. A preliminary study of charcoalified plant fossils from the Middle Jurassic Scalby Formation of North Yorkshire. Special Papers in Palaeontology, 49, 101 -111. CREBER, G.T. 1977. Tree rings: a natural data-storage system. Biological Reviews, 52, 349-383. CREBER, G.T. & CHALONER, W.G. 1984. Influence of environmental factors on the wood structure of living and fossil trees. Botanical Review, 50, 357-448. CREBER, G.T. & CHALONER, W.G. 1985. Tree growth in the Mesozoic and Early Tertiary and the reconstruction of palaeoclimates. Palaeo geography, Palaeoclimatology, Palaeoecology, 52, 35-60. DALRYMPLE, R.W., ZAITLIN, B.A. & BOYD, R. 1992. Estuarine facies models: conceptual basis and stratigraphic implications. Journal of Sedimentary Petrology, 62, 1130-1146. FALCON-LANG, HJ. 2000. The relationship between leaf longevity and growth ring markedness in modern conifer woods and its implications for palaeoclimatic studies. Palaeogeography, Palaeoclimatology Palaeoecology, 160, 317-328. FERUGLIO, E. 1950. Descripcion geologica de la Patagonia,Vol 1. Yacimientos Petroleosy Fiscales, Buenos Aires. FRANCIS, J.E. 1984. The seasonal environment of the Purbeck (upper Jurassic) fossil forests. Palaeogeography, Palaeoclimatology, Palaeoecology, 48, 285-307. FRITTS, H.C. 1976. Tree Rings and Climate. Academic Press, London. FRITTS, H.C., SMITH, D.G., CARDIS, J.W. & BUDELSKY, C.A. 1965, Tree-ring characteristics along a vegetation gradient in northern Arizona. Ecology, 46, 393-401. GOTHAN, W. 1925. Sobre Restos de Plantas Fosiles Procedentes de la Patagonia. Boletin de la Academia Nacional de Ciencias (Cordoba), 28, 197-212. GUST, D.A., BIDDLE, K.T., PHELPS, D.W. & ULIANA, M.A. 1985. Associated Middle and Late Jurassic volcanism and extension in southern South America. Tectonophysics, 116, 223-253. GULISANO, C.A. & HINTERWIMMER, G. 1986. Facies deltaicas del Jurasico medio en el oeste de Neuquen. Boletin de Informaciones Petroleras 3a EpocaIII,2-3\. JEFFERSON, T.H. 1982. Fossil forests from the lower Cretaceous of Alexander Island, Antarctica. Palaeontology, 25, 681-708.

PALAEOCLIMATE FROM BAJOCIAN WOOD JONES, T.P. 1993. New morphological and chemical evidence for a wildfire origin for fusain from comparisons with modern charcoal. Special Papers in Palaeontology, 49, 113-123. KRAUSEL, R. 1949. Die fossilen Koniferen-H01zer, 1. Palaeontographica B, 89, 83-203. LAMARCHE, V.C. 1982. Sampling strategies. In: HUGHES, M.K., KELLY, P.M., PILCHER, J.R. & LAMARCHE, V.C. (eds) Climate From Tree Rings. Cambridge University Press, Cambridge, 4. LEGARRETA, L. & ULIANA, M.A. 1991. Jurassic Cretaceous marine oscillations and geometry of a backarc basin fill, central Agentine Andes. In: MCDONALD, D.I.M. (ed.) Sedimentation, Tectonics and Eustacy. International Association of Sedimentologists, Special Publications, 12, 429-450. LEGARRETA, L. & ULIANA, M.A. 1996. The Jurassic succession in west-central Argentina: stratal patterns, sequences and paleogeographic evolution. Palaeogeography, Palaeoclimatology, Palaeoecology, 120, 303-330. MclLROY, D. 2004. Sedimentology and ichnofabrics of the He Formation, Jurassic, Offshore Mid, Norway. In: MclLROY, D. (ed.) The Application of Ichnology to Stratigraphic and Palaeoenvironmental Analysis. Geological Society, London, Special Publications, 228, 237-273. MclLROY, D., FLINT, S. & HOWELL, J. 1999. Applications of high resolution sequence stratigraphy to reservoir prediction and flow unit definition in aggradational tidal successions. In: GCSSEPM Foundation, 19th Annual Research Conference. Advanced Reservoir Characterization. SEPM, Tulsa, 121-132. MclLROY, D., FLINT, S., HOWELL, J.A. & TIMMS, N. 2005. Sidementology of the tide-dominated Jurassic Lajas Formation, Neuquen Basin. Argentina. In: VEIGA, G.D., SPALLETT, L.A., HOWELL, J.A. & SCHWARZ, E. (eds) The Nequen Basin: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 83-107. MARTINEZ, M.A., GARCIA, V.M. & QUATTROCCHIO, M.E. 1996. Analisis de componentes principales aplicado al estudio palinologico del Jurasico Medio de Cuenca Neuquina, Argentina. XIII0 Congreso Geologico Argentina y III0 Congreso de Exploracion de Hidrocarburos, 5, 171-179. MARTINIUS, A.W., KAAS, L, N^ESS, A., HELGESEN, G., KJ^REFJORD, J.M. & LEITH, D.A. 2000. Sedimentology of the heterolithic and tide-dominated Tilje Formation (Early Jurassic, Halten Terrace, offshore mid-Norway). In: MARTINSEN, O.J. & DREYER, T. (eds) Sedimentary Environments Offshore Norway - Paleozoic to Recent. Norwegian Petroleum Society, Special Publications, 9, 103-144. PARRISH, J.T. & SPICER, R.A. 1988. Middle Cretaceous woods from the Nanushuk Group, central North Slope, Alaska. Palaeontology, 31, 19-34. QUATTROCCHIO, M. & VOLKHEIMER, W. 1990. Jurassic and Lower Cretaceous dinocysts from Argentina: their biostratigraphic significance. Review of Palaeobotany and Palynology, 65, 319-330.

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QUATTROCCHIO, M., SARJEANT, W.A.S. & VOLKHEIMER, W. 1996a. Marine and Terrestrial Jurassic Microfloras of the Neuquen Basin (Argentina): Palynological Zonation. In: RICCARDI, A.C. (ed. Advances in Jurassic Research. GeoResearch Forum 1-2. Trans-Tech Publications, ZiirichUetikon, 167-178. QUATTROCCHIO, M., SARJEANT, W.A.S. & VOLKHEIMER, W. 1996/7. Paleogeographic changes during the Middle Jurassic in the southern part of the Neuquen Basin, Argentina. In: RICCARDI, A.C. (ed.) Advances in Jurassic Research. GeoResearch Forum 1-2. Trans-Tech Publications, ZiirichUetikon, 467-483. RAMOS, V.A. 1988. The tectonics of the Central Andes; 30 to 33 S Latitude. In: CLARK, S.P.J., BURCHFIEL, B.C. & SUPPE, J. (eds) Processes in Continental Lithospheric Deformation. Geological Society of America, Special Paper, 218, 31-54. REES, P.McA., ZIEGLER, A.M. & VALDES, P.J. 2000. Jurassic phytogeography and climates: New data and model comparisons. In: HUBER, B.T., MACLEOD, K.G. & WING, S.L. (eds) Warm Climates in Earth History. Cambridge University Press, Cambridge, 297-318. RICCARDI, A.C., GULISANO, C.A., MOJICA, J., PALACIOS, O., SCHBERT, C. & THOMSON, M.R.A. 1992. Western South America and Antarctica. In: WESTERMAN, G.E.G. (ed.) The Jurassic of the Circum-Pacific. World and Regional Geology 3. Cambridge University Press, Cambridge, 122161. RIGGS, E.S. 1926. Fossil hunting in Patagonia. Natural History, New York, 26, 536-544. SCHIMPER, W.P. 1869-1874. Traite de paleontologie vegetale ou la flore du monde primitif. Litterature phytopaleontologique, Paris, Plates 1-56 (1869); 2, 1-522, plates 57-84 (1870); 523-968, plates 85-94 (1872); 3, 1-896, plates 95-110 (1874). SEWARD, A.C. 1919. Fossil Plants, Volume IV. Cambridge Biological Series. Cambridge University Press, Cambridge. SPEGAZZINI, C. 1924. Coniferales fosiles Patagonicas. Annales Societiedad Cientifica Argentina, 98, 125 - 139. STIPANICIC, P.N. & BONETTI, M.I.R. 1970. Posicion estratigraficas y edades de las principales floras jurasicas argentines; II, Floras doggerianas y malmicas. Ameghiniana 7, 101-118. ULIANA, M.A. & BIDDLE, K.T. 1987. Permian to la Cenozoic evolution of northern Patagonia; main tectonic events, magmatic activity, and depositional trends. In: McKENZiE, G.D. (ed.) Gondwan Six; Structure, Tectonics, and Geophysics. American Geophysical Union, Geophysical Monograph, 40,271-286. VAN WAGONER, J.C., POSAMENTIER, H.W., MITCHUM, R.M.JR., VAIL, P.R., SARG, J.F., LOUTIT, T.S. & HARDENBOL, J. 1988. An overview of the fundamentals of sequence stratigraphy and key descriptions. In: WILGUS, C.K., HASTINGS, B.S., KENDALL, C.C.ST.C, POSAMENTIER, H.W Ross, C.A. & VAN WAGONER, J.C. (eds) Sea-lev

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WIELAND, G.R. 1935. The Cerro Cuadrado Petrified Forest. Publications of the Carnegie Institution of Washington, 449. WINDHAUSEN, W.S. 1924. Lineas generates de la con stitucion geologica de la region situada al oeste del Golfo de San Jorge. Boletin Academia Ciencifica Cordoba, 27, 167-320. ZAITLIN, B.A., DALRYMPLE, R.W. & BOYD, R. 1994 The stratigraphic organisation of incised-valley systems associated with sea-level change. In: DALRYMPLE, R.W., BOYD, R. & ZAITLIN, B.A. (eds) Incised Valley Systems: Origin and Sedimentary Sequences. SEPM, Special Publications, 57, 45-60. ZAVALA, C. 1996. High-resolution sequence stratigra phy in the Middle Jurassic Cuyo Group, South Neuquen Basin, Argentina. In: RICCARDI, A.C. (ed.) Advances in Jurassic Research. GeoResearch Forum 1-2. Trans-Tech Publications, ZiirichUetikon, 295-304. ZIEGLER, A.M., PARRISH, J.M. ETAL. 1993. Early Mesozoic phytogeography and climate. Philosophical Transactions of the Royal Society B, 341, 297-305.

Jurassic marine reptiles of the Neuquen Basin: records, faunas and their palaeobiogeographic significance ZULMA GASPARINI & MARTA FERNANDEZ Departamento Paleontologia Vertebmdos, Facultad de Ciencias Naturales y Museo, Universidad Nacional de La Plata (1900) La Plata, Argentina (e-mail: zgaspari @ museo.fcnym. unlp. edu. ar; martafer@ museo.fcnym. unlp. edu.ar) Abstract: The largest diversity of Jurassic marine reptiles from Gondwana has been recorded in the Neuquen Basin. Although the Early Jurassic records are limited, the records from the Middle Jurassic, especially the early Bajocian plesiosaurs and ichthyosaurs, are unique in the world. The highest abundance and taxonomic diversity is recorded from the Late Jurassic (Tithonian). Pleurodiran (Notoemys) and cryptodiran turtles (Neusticemys), ichthyosaurs (Ophthalmosaurus, Caypullisaurus), pliosaurs (Pliosaurus, Liopleurodori) and crocodilians (Geosaurus, Metriorhynchus, Dakosaurus) are members of this rich marine herpetofauna. Except for Notoemys, the other reptiles are pelagic. Their habits and large size (Caypullisaurus, Liopleurodon, Dakosaurus} suggest that they entered the sheltered Neuquen Basin occasionally, possibly for reproduction in protected areas. The model of a protected basin, open to the Pacific Ocean through gaps in an emergent by island-arc complex, fits well with the ecological requirements inferred for most of the pelagic reptiles. The Caribbean Corridor may also have played a significant role as a seaway at least during the Middle Jurassic and probably before. New findings in Tithonian and Berriasian sediments of the Neuquen Basin suggest that there was no massive extinction among the marine herpetofauna at the Jurassic-Cretaceous boundary.

Marine reptiles had an outstanding role as large predators within the Jurassic seas (Massare 1997). Within a food web some of them, such as the ichthyosaurs, can be defined as intermediate species, having both predators and prey within the web (Pimm et al. 1991), while pliosaurs and large metriorhynchids could be considered as top predators. Remains of Jurassic marine reptiles have been found in all continents, including Antarctica. However, the knowledge of the evolutionary history of these groups is limited when compared to that of terrestrial Mesozoic reptiles such as dinosaurs or crocodiles. This is a largely a consequence of insufficient field work and the cost of specimen preparation. The difficulties associated with systematics are also significant. The extensive, postburial modification of their skeleton in many cases makes the primary discernment of homologies difficult (i.e ichthyosaur fore fin homologies: Motani 19990; Fernandez 2001), resulting in problems for systematic studies, in this sense, it is noteworthy that the adaptation to sea life requires certain morphological and physiological changes that constrain the anatomical diversity. As the majority of documented findings are from the northern hemisphere our

knowledge of the evolutionary history of the different clades is geographically biased, as frequently the hypotheses are supported exclusively by Laurasiatic records. Marine reptiles are not a monophiletic group and, hence, the name is used informally and includes any reptile able to grow and feed in a salt-water environment (Hua & Buffetaut 1997). Jurassic marine reptiles include ichthyosaurs, pliosaurs and plesiosaurs, turtles and marine crocodiles (Massare 1988). Except for a few records from eastern Africa (Bardet & Hua 1996), the largest amount and diversity of Jurassic marine reptiles of the northern hemisphere was found in western Europe (Martill & Hudson 1991; Bardet 1995; Benton & Spencer 1995; Hauff & Hauff 1981). Most of them were discovered during the 18th, 19th and at the beginning of the 20th centuries (Buffetaut 1992; Taylor 1997). These specimens, with the original studies and recent revisions, are the basis of the present knowledge of marine reptiles of this period (Brown 1981; Taylor 1997; Motani 1999£; Sander 2000; O'Keefe 2001; McGowan & Motani 2003). There are also records from eastern Europe and Asia, but these are very incomplete and often have poorly defined

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Cas Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 279-294. 0305-8719/05/$15.000...© The Geological Society of London 2005.

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geographic and stratigraphic constraint (Storrs et al. 2001). In North America, despite the excellent Jurassic outcrops and the intensive study of their biota, the interest in prospecting for marine reptiles has been significantly low (Gilmore 1906; Bakker 1993; O'Keefe & Wahl 2003). In western Cuba a Middle-Late Oxfordian herpetofauna has been found. This fauna is unique in its diversity for that time (Fernandez & Iturralde-Vinent 2000; de la Fuente & Iturralde-Vinent 2001; Gasparini & Iturralde-Vinent 2001; Gasparini et al. 20020, 2004). Southern hemisphere occurrences are even scarcer than those in the northern hemisphere, and consist mostly of isolated remains of ichthyosaurs and plesiosaurs in Australia (Long & Cruickshank 1998), Madagascar (Fernandez 1997a), and the Antarctic Peninsula (Whitham & Doyle 1989). In South America, there are unpublished specimens in Peru, while in Chile generally incomplete marine reptiles are known from different Jurassic stages (Gasparini 1985; Gasparini et al. 2000; Shultz et al. 2003). The only exception to this scarce southern hemisphere record are the Jurassic reptiles found in the Neuquen Basin. These have been known since the 19th century, and today, as a result of continous prospecting over the last 25 years, offer the most complete record of Gondwanan marine reptiles. The Neuquen Basin, in central-western Argentina, encompasses an area of approximately 120 000 km2 (Fig. 1). The basin contains extensive outcrops of Jurassic-Albian marine deposits (Mitchum & Uliana 1985; Legarreta & Uliana 1991; Yrigoyen 1991). The structural and stratigraphic evolution of this basin has been documented by several authors (Legarreta & Gulisano 1989; Legarreta & Uliana 1996). In many of its lithostratigraphic units, especially the Jurassic ones, a rich and diverse herpetofauna has been discovered (Gasparini & Fernandez 1996, 1997; Fernandez 1997Z?, 2000). Recent systematic, palaeobiological andpalaeobiogeographic studies have been combined with interdisciplinary approaches that analysed the environments and the taxa inhabiting them. This has enlarged the knowledge of the marine Jurassic herpetofauna in general, and particularly that of the fauna that lived in or visited this basin. The aims of this paper are to synthesize the exisiting data on the basin, to compare these to the previously documented Jurassic faunas of the northern hemisphere, to consider the structure of the communities and their dispersal, and, finally, to consider the occurrence or not of an extinction event during the Tithonian-Berriasian transition.

Jurassic marine reptiles in the Neuquen Basin The history of the discoveries and associated taxonomic studies has been briefly reported in several syntheses (Gasparini 1985, 1996; Gasparini & Fernandez 1996, 1997). New data and those that concern the interpretation of the faunas and their distribution are mentioned below. Early Jurassic All the Early Jurassic reptile specimens of the Neuquen Basin are so incomplete that it is impossible to identify them below the order level. In addition, most of them have no precise geographic and stratigraphic provenance. Rusconi (1948<2, 1949) named two new species on the basis of ichthyosaur centra found at San Juan (Fig. 2a) later referred as nomen vanum (Gasparini 1985). More recently, Fernandez & Lanes (1999) reported ichthyosaur vertebrae from the northern margin of the Atuel River in southern Mendoza province (Fig. 1). These were from the Puesto Araya Formation (Volkheimer 1973), which is considered to be Lower Sinemourian (Riccardi et al. 1988; Lanes 2002). Huene (1927) described a new species of teleosaurid crocodilian, Steneosaurus gerthi (Fig. 2b) on the basis of two dorsal vertebrae found at the area of Portezuelo Ancho (Fig. 1), which he referred to the Upper Lias. The dorsal vertebrae of teleosaurids and metriorhynchids are very difficult, if not impossible, to differentiate. In addition, both families were present toward the end of the Lias, at least in Europe (Buffetaut 1982). Consequently, the Mendoza specimen is referred to Thalattosuchia indet. Middle Jurassic Marine reptiles from the Lower Bajocian and Lower Callovian have been found in central and northern Chile (Gasparini et al. 2000), and in SW Neuquen. The area with the highest diversity of taxa and most complete specimens is Chacaico Sur, Neuquen (Fig. 1), where the transition between the Los Molles and Lajas Formations is well exposed (Spalletti et al. 1994). At the lowest section of the Los Molles Lajas transition (Emileia giebeli Zone, Early Bajocian) some pliosauroid plesiosaurs have been discovered, among which an almost complete three-dimensional skull (Maresaurus coccai: Gasparini 1997) is outstanding. Maresaurus (Fig. 2c) shares characters with pliosaurs from the European Early Jurassic (Rhomaleosaurus

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Fig. 1. Neuquen Basin with of the principal localities where marine reptiles were found: (1) Paso del Espinacito; (2) Quebrada Honda, Espinacito; (3) Puesto Anaya; (4) Portezuelo Ancho; (5) Chacaico Sur; (6) Chacay Melehue; (7) Quebrada Remoredo; (8) Bardas Blandas; (9) Cajon Grande; (10) Cerro Lotena; (11) Los Catutos; (12) Arroyo Trincajuera; (13) Yesera Tromen-Pampa Tril.

Seeley) and Middle and Late Jurassic (Simolestes Andrews) from Tethys. A caudal vertebra of a Thalattosuchian crocodile was found at the same stratigraphic level as Maresaurus. This is the first record of Middle Jurassic marine crocodiles in Argentina. Recently, the presence of metriorhynchid thalattosuchians has been proved for the Early Bajocian in Chile (Gasparini et al 2000) and for the Late Bathonian at Chacay Melehue

(Fig. 1) in Neuquen province (Gasparini et al. 2005). It is significant to note that no ichthyosaurs are known for the Aalenian-Bathonian intervals outside the Neuquen Basin (Bardet 1995). Recently, Fernandez (2003) described an ophtalmosaurian ichthyosaur from the Aalenian Bajocian boundary in southern Mendoza (Fig. 1), this being the oldest record of

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Fig. 2. Early and Middle Jurassic marine reptiles from the Neuquen Basin, (a) Ichthyosaur vertebrae, Early Jurassic of San Juan, modified from Rusconi (1948&); (b) thalattosuchian vertebrae, Late Lias, Mendoza, modified from Huene (1927); (c) skull of Maresaurus coccai, early Bajocian, Chacaico Sur, modified from Gasparini (1997); (d) Ophthalmosaurian paddle, Aalenian-Bajocian boundary of Mendoza, modified from Fernandez (2003); (e) skull of Chacaicosaurus cayi, early Bajocian, Chacaico Sur, modified from Fernandez (1994); (f) skull ofMollesaurusperiallus, early Bajocian, Chacaico Sur, modified from Fernandez (1999); (g) fragment of snout of 'Stenopterygius grandis' early Bajocian, Curru Charahuilla, modified from Cabrera (1939); (h) pliosaurid vertebra, early Callovian, Chacaico Sur, modified from Gasparini & Spalletti (1993); (i) vertebrae of cf. Muraenosaurus sp., early Callovian, Chacaico Sur, modified from Gasparini & Spalletti (1993); and (j) vertebrae of cf. Cryptoclidus sp., early Callovian, Chacaico Sur, modified from Gasparini & Spalletti (1993).

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ophthalmosaurian ichthyosaurs worldwide (Fig. 2d). The first early Bajocian ichthyosaurs to be documented from the basin were found at Chacaico Sur-Charahuilla. One of these taxa is Chacaicosaurus cayi Fernandez 1994, a longirostran ichthyosaur with a proportionally small orbit (Fig. 2e) related to Stenopterygius, from the European Early Jurassic (Fernandez 1999). The other is Mollesaurus perihallus Fernandez 1999, an ophtalmosaurian with very large orbit (Fig. 2f), closely related to Ophthalmosaurus. In Charahuilla, a locality near Chacaico Sur, a rostral fragment of another ichthyosaur was found which Cabrera (1939) referred to the new species 'Stenopterygius grandis'. Given the fragmentary material, the validity of this name is still uncertain (Fig. 2g). Reptiles were also found at Chacaico Sur in levels of the Lajas Formation referred to as early Callovian (Gasparini & Spalletti 1993). Unlike the well-preserved and even articulated specimens from the early Bajocian, these are numerous isolated vertebrae and other post-cranial fragments. Based on vertebral morphology, pliosauroids (Fig. 2h) and plesiosauroids could be identified. Some vertebrae (Fig. 2i) were referred to elasmosaurids (cf. Muraenosaurus sp.) and others (Fig. 2j) to cryptoclidids (cf. Cryptoclidus sp.) (Gasparini & Spalletti 1993). Late Jurassic All the Late Jurassic marine reptiles from the Neuquen Basin were found in the Vaca Muerta Formation. The base of the formation was deposited as a consequence of a rapid transgressive event biostratigraphically dated as Tithonian Valanginian. This flooding extended over a large part of the basin (Spalletti et al. 2000). During the 1940s Carlos Rusconi described several Tithonian ichthyosaur and plesiosaur species found in Mendoza (Rusconi 1948<2, b, 1949, 1967). These taxa are not valid and a supposed plesiosaur (Purranisaurus potens Rusconi 1948b) is a Metriorhynchidae crocodile (Gasparini 1985) (Fig. 3a). However, the works of Rusconi demonstrated the amount and diversity of marine reptiles that might be discovered if a systematic survey was undertaken. His initial studies and subsequent fieldwork carried out by other researchers (mainly by the Museo de La Plata staff) have produced a collection from the Tithonian that demonstrates the highest taxonomic diversity. Remains of Tithonian reptiles have been found in many localities in Neuquen and south of Mendoza (Gasparini 1985; Gasparini et al. 1999). The four most fossiliferous localities

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represent the early Tithonian (Cerro Lotena, Fig. 1), middle Tithonian (Los Catutos, Fig. 1), late Tithonian (Arroyo Trincajuera, Fig. 1) and Tithonian-Berriasian (Yesera del TromenPampa Tril, Fig. 1). The highest number and greatest diversity of Late Jurassic marine reptiles from South America were found at Cerro Lotena, in rocks of the Vaca Muerta Formation, in the Virgatosphinctes mendozanus Zone (early Tithonian) (Leanza 1993). The fauna includes cryptodire turtles (Neusticemys neuquind) (Fernandez & de la Fuente 1993) and pleurodire turtles (Notoemys laticentralis) (de la Fuente & Fernandez 1986, 1989; Fernandez & de la Fuente 1994) (Fig. 3b). The fauna also contains ichthyosaurs, including the holotype of Caypullisaurus bonapartei Fernandez 1991 b, a specimen that must have reached approximately 7 m long and is the most complete marine reptile of the Neuquen Basin (Fig. 3c). A pliosaur, Liopleurodon sp. (Spalletti et al. 1999a), and several specimens of the Metriorhynchidae Geosaurus araucanensis Gasparini & Dellape 1976 (Fig. 3d), have also been found. The lithographic shales of the Los Catutos Member (Vaca Muerta Formation) crop out in and around the quarry at Los Catutos (Fig. 1). This member belongs to the Windhauseniceras internispinosum (Kranz) Zone, attributed to the uppermost middle Tithonian (Leanza & Zeiss 1990). Most of the reptiles that have been discovered here were transferred to the Museo Olsacher by the quarry workers (Gasparini et al. 1995). The reptiles include ichthyosaurs, one of which is an ophthalmosaurian (Gasparini 1988; Fernandez 2001) (Fig. 3e), several specimens of the turtles Notoemys laticentralis and Neusticemys neuquina (Fernandez & de la Fuente 1993), a skull of a metriorhynchid crocodile (Geosaurus sp.), a plesiosaurs tooth and two Pterodactyloidea pterosaurs (Gasparini et al. 1987; PaulinaCarabajal & Gasparini 2002) (Fig. 3f). In late Tithonian rocks of the Trincajuera Creek (Substeueroceras koeneni Zone), several specimens of ichthyosaurs were found, one of them referred to Ophthalmosaurus sp. (Fernandez 2000), also an adult specimen of Neusticemys neuquina (Fig. 3g) and a skull fragment of a Pliosauridae (Gasparini et al. 1991 a) (Fig. 3h). Finally, in the Yesera del Tromen area, near Pampa Tril there is a sucession in which shales, laminated mudstones and marl concretions of the Tithonian-Berriasian predominate. Fifteen skeletons have been exhumed from an area of 2.1 km2 (Spalletti et al. I999a, b', Gasparini et al. 2002b). The specimens are not

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Fig. 3. Late Jurassic marine reptiles from the Neuquen Basin, (a) Purranisaurus potens, late Tithonian Jurassic of Mendoza, modified from Rusconi (1948
contemporaneous as the fossil-bearing horizons belong to a condensed sequence (Spalletti et al. I999a). However, they are one of the most conspicuous offshore herpetofaunas of the Jurassic-Cretaceous transition. Members of

this herpetofauna include Caypullisaurus bonapartei, Liopleurodon sp., Dakosaurus sp. (closely related to Dakosaurus andiniensis Vignaud & Gasparini 1996) and Geosaurus araucanensis. No turtles have been recorded.

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Palaeoenvironments and behaviours Virtually all the Jurassic reptiles from the Neuquen Basin are pelagic animals. The single exception is the turtle Notoemys, which is always recorded in marine sediments despite the retention of the appendicular skeleton of a freshwater pleurodiran (de la Fuente & Fernandez 1986; Fernandez & de la Fuente 1993, 1994). Massare (1988, 1997) analysed the hydrodynamic properties and propulsion mechanisms of different marine reptiles. The reptiles of the Neuquen Basin may be grouped according to her classification (Massare 1997) into the following Baupldne (body forms): Bauplan I (Fig. 4) is characterized by a hydrodynamic profile (streamline shaped}, narrow caudal peduncle and a large expansion of the caudal fin. This morphology favours sustained swimming and active pursuit of prey (the ichthyosaurs Chacaicosaurus, Mollesaurus, Caypullisaurus and Ophthalmosaurus). Bauplan II (Fig. 4) is represented by an elongate body, a tail with a broad base and expansion of the distal segment of the tail. This favours rapid acceleration rather than sustained swimming (the crocodilians Metriorhynchus, Geosaurus and Dakosaurus). Bauplan III (Massare 1997) is here divided into Bauplan Ilia and Illb. Bauplan Ilia (Fig. 4) includes ellipsoidal bodies with both pairs of limbs elongate, a short neck and a large skull. These, more compact bodies, have a higher pursuit capability (the pliosaurs Pliosaurus and Liopleurodori). Bauplan Illb (Fig. 4) comprises others with a longer neck, in which prevails slower movements and prey capture by ambush (the plesiosauroids, cf. Muraenosaurus and cf. Cryptoclidus). Finally, Bauplan IV (Fig. 4) is characterized by compressed bodies, with bony armour, propelled by fore fins. Reptiles of this group, unlike other marine reptiles, were not acceleration specialists, and fed on low moving benthic and nektonic animals (the turtles Notoemys and Neusticemys). In a fossil assemblage it is difficult to prove the coexistence of all the organisms (Martill et al 1994). This situation is true for the Jurassic marine reptiles of the Neuquen Basin, which are frequently found in condensed sedimentary sequences (Spalletti et al. 1994, 1999a). Even when they are found in a determined biozone, there is no other evidence confirming that they coexisted, and consequently predator-prey relationships are speculative. However, when models for palaeoenvironments are compared to the Baupldne, correlation between environments and reptile morphotypes is coherent. The gradual transition between the Los Molles and Lajas formations is well exposed at Chacaico

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Sur. Spalletti et al. (1994) recognized three sections of this transition zone, on the basis of lithofacies analysis. The lower one, which corresponds to the middle-upper part of the Los Molles Formation, is composed of dark shales and marls with intercalated sandy layers. This interval bears the largest number and diversity of Bajocian marine reptiles described from Gondwana. The interval was deposited by gravitational settling in a sublittoral environment with sporadic wave action induced by storms (Spalletti et al. 1994; Spalletti in Gasparini et al. 1997b). Accordingly, the skulls of large plesiosaurs and ichthyosaurs are articulated, as are the delicate fins of the ichthyosaurs (Fernandez 1994). The diversity of taxa among ichthyosaurs may indicate a protected area where the pursued preys came for part of their reproductive cycle. The presence of ichthyosaurs and metriorhynchid crocodilians is consistent with the abundance of cephalopods (Westermann & Riccardi 1979). Thus, the belemnites that were the main food for some of the ichthyosaurs and the small marine crocodiles (Martill 1992) are numerous at the same stratigraphic levels as these reptiles at Chacaico Sur. Finally, a pelagic predator, the pliosaur Maresaurus, may have been searching for fish and other reptiles. The middle section is the upper part of the Los Molles Formation and belongs to a lowershoreface environment. The upper section (lower part of the Lajas Formation) consists of alternating bioturbated wackestones and sandstones with horizontal or low-angle crossbedding, as well as coquinas. These are typical of an upper-shoreface environment with bedload sedimentation due to high-energy wave-induced currents (Spalletti et al. 1994). In this upper section, which is lower Callovian in age, reptile bones appear loose and broken, mostly comprised of vertebral centrae and limb bone fragments (Gasparini & Spalletti 1993). The disarticulation is attributed to wave action. All the bone fragments belong to plesiosaurs, a few are pliosaurs (Bauplan Ilia) while most are long-necked plesiosaurs (Bauplan Illb), which may have had more nearshore habits (Massare 1997). The diversity of the palynoflora also suggests proximity of land (Quattrocchio et al. 1996; Martinez 2002). At the Cerro Lotena area, the palaeoenvironment of the Lower Tithonian has been interpreted as temperate-warm shallow waters (Leanza 1980, 1993). The abundance of nektonic and benthic invertebrates suggests good sun penetration in the external part of the basin, despite the strong continental supply (plant fragments and trunks). Vertebrates include bony fish and

Fig. 4. Body forms or Bauplane modified from Massare (1997), represented in the fossil record of the Neuquen Basin.

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the first Jurassic hybodontiform shark spine recorded in South America (Cione 1992; Cione et al 2002). This section also includes the largest diversity of marine reptiles in the Neuquen Basin. Thus, there were reptiles of more coastal habits (Notoemys), others which may have been not efficient sustained swimmers, such as the turtle Neusticemys and the small crocodiles Geosaurus, as well as large offshore predators and swimmers (Pliosaurus, Liopeurodon, Caypullisaurus). Palaeogeographic reconstructions of the Middle Tithonian, show that the area of Los Catutos lay 100 km NW of the SE coast of the Neuquen Basin, under subtropical conditions (Leanza & Zeiss 1990; Legarreta & Uliana 1991,1996). The lithographic shales were deposited in a warm sea with a high sedimentation rate that permitted the preservation of a diverse assemblage of organisms (Cione et al. 1987; Leanza & Zeiss 1990; Gasparini et al. 1995; Scasso & Concheyro 1999; Scasso et al. 2002), including the first ray recorded out of Europe (Cione 1999). Except for a middle-sized ichthyosaur, the remaining reptiles discovered from Los Catutos are relatively small (Notoemys, Neusticemys, Geosaurus and an ophthalmosaurian ichthyosaur). No large offshore predators have been recorded to date. In the Vaca Muerta Formation cropping out to the east of Yesera del Tromen, shales, laminated mudstones and marl concretions prevail (Spalletti et al. I999b). Palaeoenvironmentally these deposits represent the basin portion of a marine ramp, and were deposited from the settling of carbonate and siliceous particles over a predominantly anoxic seafloor (Spalletti et al. I999a, b). According to Spalletti et al. (19996, 2000) the isolation of the basin, a low surrounding relief and a mainly dry climate were the main causes of the development of a strongly anaerobic environment at the sediment-water interface of the Tithonian-Berriasian ramp. They stated that these conditions resulted in water stratification and the development of a strong thermopicnocline below storm wave base. Thus, the waters over the thermo-picnocline were well oxygenated with a diverse biota, which is recorded in the more coastal environments. In the deeper water parts of the basin, such as the Yesera del Tromen-Pampa Tril, the reptilebearing intervals are largely devoid of invertebrates (Spalletti et al. 19996; Gasparini et al. 20026). Where present, invertebrates include rare ammonites and even more scarce isolated and badly preserved bivalves. Likewise, fish remains are rare and also badly preserved. All reptiles found are offshore pelagic forms

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(Caypullisaurus, Dakosaurus, Geosaurus), lacking those of the more coastal environment, such as turtles and long-necked plesiosaurs. The reptiles skeletons, whether complete or incomplete, lie articulated with a high degree of preservation. This preservation and the position of the carcasses are related to the absence of kinetic energy in the deep waters, and the presence of a soft muddy substrate (Spalletti et al. I999a, b). Similar conditions of very soft or soupy substrates have been recognized as one of the main factors in the marine vertebrate conservation in other systems such as the Posidonian Shales (Martill 1993). The Neuquen Basin has been traditionally interpreted as a narrow seaway open to the Oriental Pacific in the area south of the presentday San Juan Province that widened in the area of Neuquen Province (Legarreta & Uliana 1996). However, the presence of so many large pelagic reptiles would be unlikely in a relatively narrow and semi-closed basin 1200 km in length. In such a semi-closed basin the salinity would have been high. Under these conditions, the biodiversity would be expected to be low. But the contrary is the case. The biodiversity was high, especially near the coast in areas such as Chacaico Sur (Volkheimer 1973; Westermann & Riccardi 1979; Spalletti et al. 1994), Cerro Lotena (Leanza 1980) and Los Catutos (Gasparini et al. 1995; Scasso et al. 2002). The salinity is interpreted to have been normal and sea-water temperatures, at least in the area of Yesera del Tromen-Pampa Tril, would have varied between 22 and 27 °C (Matheos et al. 2000). If the palaeodiversity were high and the salinity and temperature normal, then it is possible that the basin was in more direct communication with the Pacific Ocean. In the model of Spalletti et al. (2000, fig. 8) the Neuquen Basin is separated from the Pacific by an island arc with frequent gaps. Such a model is the best explanation for the diversity of large pelagic predators in the basin over millions of years. What benefits could such a protected basin have for the pelagic reptiles? Although the morphological features of the reptiles found in the Neuquen Basin (except for Notoemys) suggest that these animals spent a significant part of their life cycles in open seas, a protected basin could have had several benefits even for pelagic viviparous forms. Except for the turtles, which undoubtedly spawned on land, all marine reptiles of the Neuquen Basin were viviparous. The viviparity of ichthyosaurs is based on directed evidence such as the preservation of pregnant females (Boettcher 1990; McGowan 1991). The

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interpretation of the reproductive habits of plesiosaurs (Rothschilds & Martin 1993) and metriorhynchid crocodilians, however, are controversial. Although Hua & Buffetaut (1997) stated that they reached the land for spawning, the anatomical adaptations of these crocodilians, such as their long tail with its posterior end deflected downwards and the lengthened hind limbs, enabled them to support their body on land, and limited their capability of digging in the shore sand for spawning, suggesting that they gave birth to living young (Gasparini 1978). Although no pregnant females have been discovered, juveniles and newborns have been found in several localities of the Neuquen Basin (Cerro Lotena, Los Catutos, Yesera del Tromen-Pampa Tril). This suggests that the basin might have been used for different stages of reproduction. In addition, the abundance of newborns and juveniles would have been an item in the diet of the top predators, such as Liopleurodon and Dakosaurus, found in the same stratigraphic levels.

Palaeobiogeographic distribution The Jurassic marine herpetofauna of the Neuquen Basin shows a close relationships with that of western Tethys (Fig. 5). These affinities were first recognized in the 1970s by Gasparini (1978) and Gasparini & Dellape (1976). In the following section we compare the reptiles from the different stratigraphic intervals within the basin to their European (Tethyan) counterparts. This discussion has implications for the existence of a Caribbean corridor connecting the Tethys with the Pacific (Damborenea & Mancenido 1979). The record of marine reptiles from the Early Jurassic of South America is very scarce and currently no specimens have been determined at family, genus or species level. Fragments of marine crocodiles have been found in the Early Jurassic of Chile (Huene 1927; Chong & Gasparini 1972). At this time, teleosaurids and probably metriorhynchids lived in the European Tethys (Buffetaut 1982; Vignaud 1995). Diverse evidence, mainly from invertebrates and microfossils, suggest the sporadic presence of the Caribbean Corridor (Damborenea & Mancenido 1979; Riccardi 1991; Boomer & Ballent 1996). This corridor may have played a main role in the dispersion of fish and reptiles, particularly crocodiles (Gasparini 1981, 1992) (Fig. 6a). In the Middle Jurassic, the reptiles of the Neuquen Basin fill a gap of approximately 10 Ma in which, except for the European crocodiles, the global record is very poor (Bardet 1995).

The first ophtalmosaurians are recorded in the Neuquen Basin (Aalenian-Bajocian), a clade that persisted until the Early Cretaceous with a wide geographic distribution (Motani 1999Z?; Fernandez & Iturralde-Vinent 2000; Sander 2000; Fernandez 2003). Other South American marine reptiles that showed strong affinities with those of the European Tethys include the pliosaur Maresaurus (Bajocian), a predator that shared characteristics with Rhomaleosaurus (Lias) and Simolestes (Callovian-Tithonian) of Europe (Gasparini 1997), and Metriorhynchus, in the Bajocian of Chile and Bathonian of Argentina, a genus that prevailed in the Callovian record of Europe (Gasparini et al 2000, 2005). Ophtalmosaurians (Mollesaurus periallus) and a probable stenopterygian were also present. In Europe stenopterygians are Liassic, while the most ancient ophtalmosaurians are Callovian. Finally, in early Callovian rocks, both in Chile and Argentina, there are records of taxa with very close affinities to forms that lived at this time at least in the seas that covered part of Europe. Such is the case of Metriorhynchus casamiquelai related to M. brachyrhynchus (Gasparini et al. 2000), and the long-necked plesiosurs cf. Muraenosaurus and cf. Cryptoclidus (Gasparini & Spalletti 1993). On the basis of the taxonomic affinities among crocodiles, Gasparini (1978, 1981, 1996) proposed that the most parsimonious way for the dispersal of these reptiles had to be the Caribbean Corridor (Hispanic Corridor) (Fig. 6b). While the existence of the Caribbean Corridor during the Middle Jurassic is still controversial (Iturralde-Vinent 2003), there is a lot of palaeontological evidence supporting at least a periodically discontinous seaway (Damborenea 2000). There is positive evidence supporting the opening of the Caribbean Corridor by the Oxfordian (Iturralde-Vinent 2003), as the palaeontological record of the Jagua Formation of Cuba includes ophtalmosaurian ichthyosaurs (Fernandez & Iturralde-Vinent 2000), pleurodiran turtles related to Notoemys (de la Fuente & Iturralde-Vinent 2001), Geosaurus, a metriorhynchid frequently found in the Tithonian of Europe, Mexico and Argentina (Gasparini & Iturralde-Vinent 2001), and criptoclidid plesiosaurs (Gasparini et al. 20020). The reptiles found in the Cuban Oxfordian mark the importance of the Caribean Corridor (Fig. 6c) in the reptile dispersal, and its study within a phylogenetic frame suggests that the contact between the eastern Pacific and western Tethys faunas must have occurred at least in the Middle Jurassic (Fig. 5). In the Tithonian of the Neuquen Basin, there are records of metriorhynchid crocodiles

Fig. 5. Cladograms showing phylogenetic affinities of some of the marine reptiles mentioned in the text, (a) Plesiosauroids, modified from Gasparini et al. (2002a); (b) metriorhynchids; (c) pleurodiran turtles, modified from de la Fuente & Iturralde-Vinent (2001); and (d) Ichthyosaurs, modified from Motani (\999a). Abbreviations: CC, Caribbean Corridor; EP, eastern Pacific; WT, western Tethys.

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Fig. 6. Palaeogeographic reconstruction, (a) Early Jurassic; (b) Middle Jurassic; (c) Late Jurassic. Abbreviations: CC, Caribbean Corridor; NB, Neuquen Basin; WS, Weddell Sea. Modified from Smith et al (1994).

(Geosaurus, Metriorhynchus, Dakosaurus), pliosaurians (Pliosaurus, Liopleurodori) and ichthyosaurs (Ophthalmosaurus), which were also recorded in the Late Jurassic of Europe (Bardet 1995; Bardet et al. 1997). The ichthyosaur Caypullisaurus and the turtle Notoemys are phylogenetically related to European forms. The interchange through the Caribbean Corridor probably continued, as is testified by a specimen referred to as Geosaurus sp. from Tithonian rocks of SW Mexico (Stinnesbeck et al. 1993). By the end of the Jurassic, however, active tectonism opened another passage, the Weddell Sea (Franzese et al. 2003), which separated SW Gondwana (South America-Africa) (Fig. 6c). Marine fish and ichthyosaur remains were found in the NE of the Antarctic Peninsula in late Tithonian or early Berriasian rocks (Whitham & Doyle 1989). These events offered new dispersion paths for the large marine predators. Conclusions The best quantitative and qualitative record of Jurassic marine reptiles of Gondwana belongs to the Neuquen Basin. The early Jurassic record is scarce due to the lack of systematic searches. However, the first discoveries of early Bajocian, late Bathonian and early Callovian reptiles suggest diversity of ichthyosaurs (Stenopterygidae: Chacaicosaurus cayi, Ophthalmosauridae: Mollesaurus perihallus), plesiosaurs (Pliosauroidea: Maresaurus coccai', Plesiosauroidea: cf. Muraenosaurus sp.; cf. Cryptoclidus sp.), and at least the presence of Metriohynchus sp. All the cited taxa are related to others of the middle Jurassic recorded in the European Tethys. The richest record in marine reptiles of the Neuquen Basin is that of the TithonianBerriasian, with the presence of large

ichthyosaurs (Caypullisaurus bonapartei), pliosaurs (Liopleurodon sp.), crocodilians (Metriorhynchus potens, Geosaurus araucanensis, Dakosaurus andiniensis) and turtles (Notoemys laticentralis, Neusticemys neuquind). Virtually all the Jurassic reptiles from the Neuquen Basin are pelagic, and represent all the Bauplane (body forms) recognized in the European Jurassic marine reptiles. The single exception is the turtle Notoemys, which is always recorded in marine sediments despite the retention of the apendicular skeleton of a freshwater pleurodiran. The Neuquen Basin reptiles are frequently found in condensed sedimentary sequences and, hence, no evidence confirm that they coexisted even in a certain biozone. When models for palaeoenvironments are compared to the Bauplane, correlation between environments and reptiles morphotypes is coherent. Except for the turtles, which undoubtedly spawned on land, all marine reptiles of the Neuquen Basin were viviparous. Although no pregnant female have been discovered, juveniles and newborns have been found in several localities in the Neuquen Basin. This suggests that the basin might have been used for different stages of reproduction. The model of the Neuquen Basin separated from the Pacific by an island arc with frequent gaps (Spalletti et al. 2000) is the best explanation for the diversity of large pelagic predators in the basin over millions of years. The Jurassic marine herpetofauna of the Neuquen Basin shows close relationships with that of the western Tethys, mainly among the late Jurassic forms. Towards the end of the Jurassic the sea level in the European basin dropped, resulting in less attractive habits for large offshore predators (Bardet 1995). However, the Neuquen Basin remained partially open to the

JURASSIC MARINE REPTILES Pacific during the Jurassic-Cretaceous transition, and, coincidentally, the record proves the survival of large pelagic swimmers at least in western Gondwana. We are very grateful to the editors for inviting us to contribute to this volume. For critical reviews we especially thank N. Bardet (Museum National d'Histoire Naturelle) and B. Nicholls (Royal Tyrrell Museum). C. Deschamps

helped with English translation. Research for this chapter was partially supported by Consejo Nacional de

investigaciones Cientificas y Tecnologicas (PIP 6298); Agencia de Promocion Cientifica y Tecnologica, Argen-

tina (PICT 8439) and by the National Geographic Society (grant 6882-00).)00

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Lithofacies, palaeoecology and palaeoenvironments of the Agrio Formation, Lower Cretaceous of the Neuquen Basin, Argentina DARIO G. LAZO, MARCELA CICHOWOLSKI, DEBORA L. RODRIGUEZ & M. BEATRIZ AGUIRRE-URRETA Laboratorio de Bioestratigrafia de Alta Resolution, Departamento de Ciencias Geologicas, Universidad de Buenos Aires, Pabellon II, Ciudad Universitaria, 1428 Buenos Aires, Argentina (e-mails: [email protected]; [email protected]; [email protected]; [email protected]) Abstract: The lithofacies and macrofossil guilds of the Agrio Formation (Upper Valanginian-Lower Barremian) have been analysed using evidence from sedimentological, taphonomic and palaeoecological studies. The study area is Agua de la Mula and adjacent regions in central Neuquen. Seven lithofacies have been recognized in the field, which indicate that the Agrio Formation was deposited in an open-marine, ramp depositional system under storm influence. Lithofacies indicate conditions that range from low-energy basin to high-energy inner ramp. Outer and mid-ramp deposits are the most abundant. Macrofossils have been grouped into 16 guilds based on tiering, life habit and feeding category. The guilds indicate normal benthic oxygen level, normal salinity, and soft—firm muddy and sandy bottoms. Suspension-feeders are more common than deposit-feeders suggesting the predominance of suspended food particles over deposited food resources. A low input of siliciclastics and, possibly, other palaeoceanographic conditions allowed the development of oolitic facies in the inner ramp and coral patch reefs in the upper mid-ramp for a limited period of time.

The Cretaceous marine deposits of the Neuquen Basin are richly fossiliferous; the invertebrate faunas, represented mostly by molluscs, have been studied since the 19th century. Most of these early works are of systematic character, including the important monograph of Weaver (1931). The aim of the present multidisciplinary study is to provide a comprehensive account of the lithofacies, palaeoecology and palaeoenvironments of the Agrio Formation (Lower Valanginian—Lower Barremian). This paper is based on ongoing research on the systematics, palaeoecology, and taphonomy of molluscs (bivalves, gastropods, nautiloids and ammonoids), corals, sponges, bryozoans, annelids, decapod crustaceans and echinoids. Some information on the various marine vertebrates has also been included to complement the picture of the different assemblages. Vertical ranges of fossils are accurately dated by the ammonoid zonation proposed by Aguirre-Urreta & Rawson (1997) and Aguirre-Urreta et al (2005, Fig. 1). Seven lithofacies and 16 guilds of macroinvertebrates are described and interpreted. Based on these data, three different fossil assemblages

have been used to reconstruct the distinct assemblages that occupied the Cretaceous sea during the deposition of the Agrio Formation. The Agrio Formation crops out along the eastern foothills of the Andes, from central Mendoza to southern Neuquen. This study is based on several sections located in the central part of the Neuquen Embayment, and especially in Agua de la Mula (see fig. 2 of Aguirre-Urreta et al. 2005). Owing to the high quality of the exposures and the gentle dip of the beds, the locality has become a key area for the study of palaeontology and sedimentology within the Agrio Formation.

Lithofacies of the Agrio Formation The Agua de la Mula section was analysed in detail to characterize the different lithologies, and their main palaeontological and taphonomic features. Lithofacies and sedimentary cycles were described in the field on a bed-by-bed basis taking into account the geometry, lithology and sedimentary structures. Detailed work was confined to the marine portions that comprise the majority of the succession.

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Case Study in Sequence Stratigraphy and Basin Dynamics, Geological Society, London, Special Publications, 252, 295-315. 0305-8719/05/$15.00 © The Geological Society of London 2005.

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Fig. 1. Stratigraphic log of the Pilmatue or Lower Member (Upper Valangini an-Lower Hauterivian) of the Agrio Formation in Agua de la Mula, showing the lithofacies distribution. Ammonoid zonation from Aguirre-Urreta & Rawson (1997) and Aguirre-Urreta et al (2005). Key is given in Figure 2.

GUILDS IN THE LOWER CRETACEOUS

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Fig. 2. Stratigraphic log of the Agua de la Mula or Upper Member (Lower Hauterivian-Lower Barremian) of the Agrio Formation in Agua de la Mula, showing the lithofacies distribution. Note that lithofacies 1, 6 and 7 are restricted to this member. For explanations see text and Figure 1.

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In the Agua de la Mula section the Agrio Formation consists of the three members originally described by Weaver (1931). The Lower or Pilmatue Member is 689 m in thickness and conformably overlies the Mulichinco Formation (flu vial-proximal marine siliciclastic deposits: Zavala 2000). The Middle or Avile Member, 45 m thick, is composed of aeolian, lacustrine and fluvial deposits that overlie a sequence boundary and are capped by a major flooding surface that represents a return to open-marine conditions (Veiga et al 2002) (Fig. 1). It has been interpreted as a lowstand wedge produced by a major fall in relative sea level during the middle Hauterivian (Legarreta & Gulisano 1989). The Upper or Agua de la Mula Member is 530 m thick (Fig. 2) and is capped by a further sequence boundary at the base of the overlying Huitrin Formation, which is comprised of restricted marine, fluvial and aeolian deposits (Vergani et al 2002). Lithofacies in the Pilmatue and Agua de la Mula members of the Agrio Formation typically occur within shallowing-upwards successions, which reflect a supply-dominated regime where the rate of sediment supply exceeded the rate of accommodation-space creation and sediment dispersal. Sediments are preserved as regressive deposits, which represent parasequences and parasequences sets (Swift et al 1991). The Agrio Formation has been interpreted as a storm-dominated shallow-marine environment, with mixed siliciclastic and carbonate sedimentation (Brinkmann 1994; Spalletti et al. 2001). The system represents a homoclinal ramp rather than a shelf, because the high-energy facies pass into deep-water mudstones that lack significant sediment gravity-flow deposits or slumping structures.

Lithofacies 1: black shales ('Spitidiscus' shales) Description. Lithofacies 1 is only recorded at the base of the Agua de la Mula Member. It is composed by thin laminated black shales and silty shales 10m thick, occurring in coarseningupwards cycles of 0.7-1 m. It grades upwards into lithofacies 2. This sandstone is topped by a thin hard siltstone with poor shell-packing and well-sorted fossils. Macrofauna in the black shales is represented by gastropods, shallow infaunal suspension-feeding bivalves and the ammonite Spitidiscus riccardii Leanza & Wiedmann (Fig. 2). Benthic elements are characterized by their small body size (less than 2 cm) and low diversity of taxa. They are usually

concentrated into pavements of disperse to loose shell packing, intercalated in shales. The fauna is preserved with shell, and the degrees of fragmentation and abrasion are very low. Bivalves are articulated. Signs of encrustation and bioerosion are absent. Interpretation. Lithofacies 1 was deposited in the transgressive event associated with the flooding of the continental deposits of the Avile Member. Thin laminated shales are suspension deposits of a very quiet environment deposited below the storm wave base. Well-preserved fossils and articulated shells indicate lowenergy levels, a low degree of lateral transport and rather episodic high sedimentation rates. Winnowing of fine sediments could have concentrated autochthonous benthic elements into pavements. The dark colour indicates high organic content that, in turn, is typical of dysoxic or anoxic seafloors. The presence of a low-diversity benthic association of small-sized shells also indicates an oxygen-controlled substrate. Absence of encrusting organisms (very common in other lithofacies) also indicates a low-oxygen environment (Wignall 1993). The development of dysoxia can be linked to a deeper stratified watermass. The presence of ammonites indicates normal marine salinity. Lithofacies 2: shale-dominated beds Description. Lithofacies 2 is the most common and thickest lithofacies in the Agrio Formation. It comprises olive-grey and dark-grey shales, and massive mudstones (silty claystones and siltstones), ranging from 2 to 60 m in thickness, alternating with thin coquinas, siltstones or fine-grained sandstones. The lithofacies occurs at the base of most shallowing-upwards cycles. It grades into lithofacies 3, 4, 5 or 6. These fine-grained sediments dominate both marine members. Levels of cm-sized calcareous nodules are commonly present. The nodules occasionally show borings and encrusting oysters. Fine bioturbation and concretionary thalassinoid trace fossils occur in the shales and mudstones. Intercalations of sandstones and siltstones are lensoid, up to 0.2 m in thickness. They may show sharp or erosional contacts, a basal lag of nodules or shells, normal grading, massive or plane lamination, and wave or combined flow ripples at the top. Firm substrate trace fossils are recorded at the base of the intercalations. Vertical and oblique burrows are filled with shell debris and small articulated infaunal bivalves. They penetrate into the underlying shales by up to 0.7 m. Coquinas occur as

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pavements and lenses up to 0.5 m thick, with loose to dense shell packing and normal graded sandy matrix. Parallel-orientated fossils dominate. Breakage and abrasion are minor and bivalves are commonly articulated, although some compaction or dissolution has occurred. Serpulids and oysters usually encrust external surfaces of large valves. Macroinvertebrates are represented by a diverse marine benthic fauna and ammonites preserved in calcareous nodules or variably concentrated within the shales. Intercalations of coquinas, siltstones and sandstones have reworked marine benthic fossils, nautiloids and ammonites. Mixed quality of preservation and different orientations characterize these deposits. Interpretation. Lithofacies 2 is interpreted as fair-weather suspension deposits of the outer ramp, as suggested by the abundance and thickness of shale, with uncommon and distal storm beds. Firm substrates were developed after seafloor erosion during storm periods as indicated by the presence of burrows with sharp walls and a distinct sedimentary fill. Nodular levels indicate episodic pauses of sedimentation that could enhance the origin of firm substrates. Shell beds are the result of the combination of bottom colonization, nektonic rain and concentration by winnowing of fine-grained sediments during fair-weather conditions. The seafloor was oxygenated and bottom colonization probably took place as the substrate changed consistency from a soupy to soft or even firm bottom (Wignall 1993). Shells were reworked by unusual storms and deposited distally as the storm flows waned on the muddy outer-ramp bottom. Storm events were partially erosional and mixed benthic organisms from muddy and sandy environments. Lithofacies 3: interbedded sandstone-shale Description. This lithofacies is comprised of alternating olive-grey mudstones and yellowish fine- to medium-grained sandstones or coquinas forming lenticular, wavy and flaser bedding. It ranges between 1.4 and 33.5 m in thickness, and grades into lithofacies 4, 5 or 6. It occurs in coarsening- and thickening-upwards trends and from lenticular and wavy to flaser bedding. The lithology of the mudstones varies between silty clay stones and sandy siltstones. They are planar laminated or massive with fine bioturbation. Levels of cm-sized calcareous nodules and concretionary thalassinoid trace fossils are usually present.

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The sandstone beds and coquinas have a sharp-erosional base and range between 5 and 30 cm in thickness. They show lags of reworked nodules, crudely graded bedding of matrix or bioclasts, valve stacking, massive or plane lamination, hummocky cross-stratification, ripple cross-lamination and wave ripples. Most of them are topped by a diverse trace fossil assemblage belonging to the Cruziana ichnofacies. The coquinas exhibit loose to dense shell packing and include variably preserved fossils, which are typically parallel orientated. Background shell debris and whole valves are usually mixed. Shells record variable taphonomic modifications. Benthic elements, mostly deep, shallow infaunal and semi-infaunal bivalves, occur in mudstones while coquinas have reworked bivalves, gastropods, nautiloids and ammonites. Interpretation. Lithofacies 3 is interpreted to have been deposited in a mid-ramp setting. It is transitional in nearly all aspects between lithofacies 2 and lithofacies 4 and 5. During fairweather conditions fall-out deposition occurred. This alternated with storm reworking and deposition (Aigner 1982), as seen from the interbedding of the mudstones with stacked amalgamated, graded tempestites. Autochthonous macrofauna represents colonization of the seafloor during fair-weather conditions, while reworked remains suffered variable offshore transport and, finally, deposition in the tempestites. Lithofacies 4: hummocky cross-stratified sandstone bodies Description. This lithofacies is composed of light-grey fine- to medium-grained sandstones and planar-laminated mudstones. The sandstones bodies are continuous along the entire outcrop and are easily correlated between adjacent localities, but individual beds are lens-shaped (tens-hundreds of metres long). They range between 1 and 17.5 m in thickness. Internal sedimentary structures vary considerably. Lower boundaries are generally sharp or erosional on lithofacies 2 or 3. Upper boundaries are also usually sharp or erosional under lithofacies 2, 3, 5 or 7. The sandstones bodies commonly show amalgamation of individual sandstone beds, which is indicated by the partial absence of muddy tops. Sedimentary structures within the hummocky sandstone bodies include planar lamination to low-angle cross-lamination, symmetrical to asymmetrical ripples, and ripple cross-lamination and hummocky

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cross-stratification. Sharp- to erosional-based beds, 7-50 cm thick, occur amalgamated or encased in mudstones. Normal graded, hummocky or plane-laminated sandstones are overlain by ripple cross-laminated sandstones or massive sandstones. Hummocks and swales are 20-100 cm long. In a few examples the sandstones show flute casts, convolute bedding and parting lineations. Ripples are straight-sinuous crested and interference ripples are common. Ripple crests are oriented both N-S and E-W. Burrowing is locally abundant. Densely and loosely packed coquinas alternate with sandstones. They are thin beds composed of wellsorted oyster shell debris or bimodally sorted. These shell beds are typically separated by dark-grey to black mudstones a few centimetres thick. The thickness of these mudstones can vary laterally due to erosion. The mudstones are laminated and commonly contain calcareous nodules and thin lenses of fine-grained sandstone. With the exception of some intercalated shell-rich beds, fossils are restricted to the sandstones. Macrofossils are represented by deep, shallow and semi-infaunal bivalves usually preserved in life position and shells of ammonoids and nautiloids. Interpretation. Lithofacies 4 is interpreted as having been deposited in a mid-ramp setting, but in a more proximal and probably shallower location than lithofacies 3 close to fair-weather wave base. This is evidenced by coarser sediments, larger hummocks and swales, and a higher degree of bed amalgamation. The abundance and thickness of sandstones indicate a large input of siliciclastic material into the shallow-marine system. The vertical succession of sedimentary structures within individual beds indicates storm-induced deposition and each sandstone body records the amalgamation of several storm events. Sedimentary structures in the sandstone bodies were produced by high-energy wave-generated oscillatory and unidirectional currents and combined flows (Midtgaard 1996). Benthic elements, preserved in life position, record the colonization of welloxygenated sandy bottoms. The presence of fully articulated bivalves implies rapid deposition. Parautochthonous fossils were reworked by storm processes and deposited as coquinas. These coquinas have fossils with mixed taphonomic indexes suggesting that shells were subject to long-term reworking by fair-weather waves (highly modified shells and debris), and episodic reworking and rapid deposition by storm waves and currents (complete and/or articulated shells).

Lithofacies 5: composite coquinas Description. This lithofacies is composed of densely and loosely packed coquinas, and ranges between 0.5 and 3.5 m in thickness. It records the accretion and amalgamation of up to seven individual coquinas. Beds are continuous laterally and can be easily correlated between adjacent sections, but individual shell beds are lens-shaped and can be followed only for tens of metres. Lower boundaries erosionally cut into lithofacies 2, 3 or 4. Upper boundaries are usually sharp and are overlain by lithofacies 2 or 3. The palaeontological and taphonomic features of lithofacies 5 vary considerably. Individual coquinas range between 0.1 and 0.5 m in thickness. They have an undulating erosional base often overlain by small calcareous nodules, infaunal bivalves or bone fragments. The matrix is graded or massive very-fine- to medium-grained quartz arenite, but calcitecemented biosparites also occur. Thin mudstone beds are sometimes preserved separating individual coquinas. Some shell beds have planar lamination at the base and/or wave ripples at the top. Hummocky cross-stratified shell debris is occasionally recorded. Most coquinas are a mixture of shells with different degrees of physical, biological or chemical alteration. The fossil assemblages consist of diverse marine benthic elements, shells of nautiloids and ammonoids, and remains of fishes and marine reptiles. The dominant bioclasts are bivalves, but locally gregarious serpulids or plates of echinoderms predominate. Shells are either filled with matrix material or calcite cement. At first sight shells seem randomly oriented in cross-section, but concordant (shells convex-down and convex-up), nested, and perpendicular orientations of shells and fragments locally dominate. Size sorting is usually bimodal with large bioclasts (e.g. fully articulated bivalves) loosely packed or dispersed among finer shell debris. The degrees of bioerosion and encrustation are moderate-low, and occur especially in thick-shelled bivalves and cephalopods. Interpretation. Lithofacies 5 is interpreted to have been deposited in a mid-ramp setting during reduced input of siliciclastic material. It is equivalent to the lithofacies 4, as it records the stacking and amalgamation of multiple storm events. The succession of sedimentary structures, erosional base of the coquinas and the concordant orientation of bioclasts support an origin via hydraulic concentration by storms. The presence of articulated bivalves and

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calcite-filled shells implies rapid deposition. Benthic elements were frequently reworked and deposited rapidly mixed with finer shell debris and nektonic cephalopods. Variable degrees of fragmentation, abrasion, encrustation and bioerosion are probably related to the time of residence of bioclasts on the seafloor. Lithofacies 6: oolitic and bioclastic carbonates Description. This lithofacies is restricted to the top of the Agua de la Mula Member of the Agrio Formation. It is composed of amalgamated beds of oosparite and biosparite, 0.5-7.55 m thick. Lower boundaries are erosionally cut into lithofacies 2 or 3, and upper boundaries are sharp under lithofacies 2. Individual beds are massive and lens-shaped, up to 30 cm thick. In the oosparites densely packed ooids and dispersed bioclasts are embedded in a sparry calcite cement. Biosparites consist of densely packed skeletal grains, dispersed ooids and isolated quartz grains. Ooids are mainly single, 0.15-0.5 mm in diameter, have a radial and tangential structure, and some have micritic texture. The nucleus may be a fragment of bivalve, echinoderm or a quartz grain. Skeletal grains include fragments of bivalves and echinoderm plates. Interpretation. Lithofacies 6 is interpreted as the product of deposition in an inner-ramp setting during times of reduced siliciclastic input. Inner-ramp deposition is suggested by the abundance of ooids, which are best developed in very shallow agitated warm waters (Scoffin 1987). Alternation of low- and highwater agitation is indicated by the combination of radial and tangential textures of ooids (Richter 1983). Lithofacies 7: coral patch reef Description. Lithofacies 7 occurs only near the top of the Agua de la Mula Member (Fig. 2). The level is 5 m thick and can be followed laterally for hundreds of metres. It rests on mid-ramp siliciclastic deposits of lithofacies 4 and grades upwards into very-shallow-water deposits of lithofacies 6. The coral facies rests on a thin and hard oyster shell bed that is located on top of lithofacies 4. It is characterized by a combination of the following: (1) flat and globose coral colonies in growth position; (2) reworked globose, conical and ramose coral colonies, benthic macrofauna and shell debris filling the

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spaces between the coral heads; and (3) amalgamated lensoid shell beds. Flat and globose coral colonies occur mostly in growth position and are usually very closely spaced. Flat colonies grew at the base of coral levels and are usually encrusted by ramose and globose forms. They have a large resting area, many small attachment areas and form thin sheets. Globose colonies have a large attachment area, large resting area and hemisphericalglobose shapes. They range from small colonies to large ones reaching up to 50 cm in length and 70 cm in height. Conical colonies have a large attachment area and a conical upper surface. They are smaller, reaching up to 10 cm in length and 20 cm in height. Ramose colonies have one or two main branches. They are always reworked and orientated parallel to the bedding. Branches range from a few centimetres to 1 m in height. They can grow up from flat or globose colonies. The corals are generally well preserved but recrystallized. They exhibit moderate-high degrees of micro- and macrobioerosion. The most common boring is Gastrochaenolites isp. The coral heads show low-moderate degrees of encrustation by sponges, serpulids and oysters. Pockets of shell debris and interstitial shell beds with reworked macrofauna are recorded between large coral colonies. Shell beds are up to 30 cm thick, have a sharp basal contact, close shell packing and a sparry calcite cement. Amalgamation of individual beds is common. Reworked erect bryozoans are locally abundant. The colonies are arborescent up to a few centimetres in length and 5 mm in width. Macrofauna is also represented by reworked epifaunal and shallow infaunal bivalves, regular echinoids and gastropods. Corals in life position are capped by amalgamated sets of shell beds up to 1.60 m thick. Each shell bed is lens-shaped and ranges in thickness from 15 to 30 cm with an ero sional base. Internally, they display massive bedding and close shell packing of well-sorted bioclasts (up to 4 cm). Moderate-high levels of fragmentation and abrasion of bioclasts are common. Skeletal components include fragments of massive and ramose corals, erect bryozoans, spines and plates of echinoids, and epifaunal bivalves. Interpretation. Corals probably started to grow after a deepening interval on sandy or bioclasticrich substrates during a period of time of reduced siliciclastic input. As no true hardgrounds were recorded, the corals may have managed to encrust a firm sediment, which is shown by the predominantly irregular, small attachment

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surfaces. Apart from this occurrence at the top of the succession, coral build-ups are apparently absent. The coral association represents a patch-reef structure developed in a shallowingupwards mid-ramp setting with frequent storm reworking. High degrees of fragmentation and abrasion, well-sorted bioclasts and absence of fine-grained sediments point to a location near fair-weather wave base. Storm influence is inferred from interstitial shell beds located between the coral heads and amalgamated sets of thin shell beds located on top of coral heads. Build-ups are laterally discontinuous and do not represent a barrier system. Overall, high siliciclastic input from the continent and lack of true hardgrounds seem to be possible reasons for the lack of coral reefs in the Agua de la Mula Member.

Guild analysis 'Guild' is a functional or ecological category that groups species having similar exploitation patterns. This term is useful in the comparative study of fossil associations as the same guild may be represented in several different palaeocommunities without regard to taxonomic groups (Root 1967). The term guild is here used as a combination of the following: (a) tiering (vertical partitioning): endobenthic (shallow or deep), semi-endobenthic, epibenthic, nekto-benthic and nektonic; (b) life habit (main behaviour): burrower, borer, recliner, mobile, attached (byssus, cement and/or roots) and swimmer; and (c) feeding category (trophic group): suspension-feeder, deposit/detritusfeeder, browser, carnivore, scavenger and predator. Thus, each guild is not restricted to a given structural plan of the body, as, for example, the guild 'epibenthic cemented suspension-feeders' is represented by bivalves, bryozoans and serpulids. In addition, specific examples and their main taphonomic features are added in each guild recorded in the Agrio Formation (Fig. 3). Endobenthic Shallow-burrowing suspension-feeders. This guild is composed of burrowing bivalves and gastropods that lived at least partially buried. Five bivalve superfamilies are included in this guild (Fig. 3). Specific examples are Cucullaea gabrielis Deshayes in Leymerie, Trigonia carinata Agassiz, Steinmanella pehuenmapuensis Leanza, S. transitoria (Steinmann), Eriphyla argentina Burckhardt, Disparilia sp. and Ptychomya koeneni Behrendsen. C. gabrielis (Arcoidea, Cucullaeidae) and Steinmanella species

(Myophorelloidea, Myophorellidae) are conspicuous and easy to recognize elements within the whole Agrio Formation. They occur as pristine articulated shells in situ in low-energy muddy bottoms, and reworked with variable taphonomic signature in sandstones and coquinas (Lazo 2004). P. koeneni (Crassatelloidea, Ptychomidae) has divaricate ribs that are transversally asymmetric and blade shape. It is interpreted as one the fastest burrowers of the bivalve assemblages of the Agrio Formation. Protohemichenopus neuquensis Camacho is a prominently alate aporrhaid gastropod (Prosobranchia, Mesogastropoda, Stromboidea) recorded in the Lower and Upper members of the Agrio Formation. It usually appears in shales and mudstones, but reworked shells in coquinas are also occasionally recorded. The living species Aporrhais pespelicani (Linne) is a shallow infaunal suspension-feeder that produces two tubes in sand for water circulation one anterior and incurrent, and the other posterior and excurrent. As the water flows in the anterior cavity the plankton sticks to mucus and then is ingested by the mobile proboscis (Schafer 1972). Deep-burrowing suspension-feeders. This guild is composed of deep-burrowing bivalves (more than 3 cm in depth). Although connected to the interface, these species can occupy the deepest tier within the sediment. Adults may reach a length of 12.2 cm with a very wide siphonal gape, suggesting a burrowing depth of about 25-30 cm. Three bivalve superfamilies are included in this guild (Fig. 3). Specific examples are Aphrodina (A) quintucoensis (Weaver), Panopea gurgitis (Brongniart) and Pholadomya gigantea (J. de C. Sowerby). P. gurgitis (Hiatelloidea, Hiatellidae) has a very thin and finely ribbed shell, which is gaped at the posterior margin. It is very common throughout the sequence and is locally abundant. It is recorded as articulated shells or moulds in mudstones, sandstones and coquinas. Specimens in life position are frequently observed on top levels of lithofacies 4. P. gigantea (Pholadomyiodea, Pholadomyidae) has a thin shell with asymmetric radial ribs, cylindrical shape and siphonal gape. This species is recorded scarcely in coquinas along the Agrio Formation and never peaks in abundance. Burrowing suspension-feeders with symbiotic bacteria. This guild is composed only of Sphaera koeneni (Behrendsen), which is a fimbriid bivalve (Veneroida, Lucinoidea). It is recorded mainly in the Pilmatue Member of the

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Fig. 3. The 16 macrofossil guilds recognized in the Agrio Formation, including the taxa belonging to each one. Note that drawings are not to scale.

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Agrio Formation but it is not abundant. It occurs usually in calcareous nodules in the lithofacies 2. Based on Recent relatives, these animals probably burrowed deeply with their hinge uppermost, maintaining an anterior, mucus-lined inhalant connection to the surface formed by the highly extensible cylindrical foot. Boring suspension-feeders. This guild consists of boring bivalves and endolithic trace fossils produced on thick-shelled bivalves, coral colonies and calcareous nodules. Boring bivalves are frequently found articulated within their perforations, but borings sometimes appear filled only with sediment or calcareous cement. In the Agrio Formation boring bivalves are represented only by the Family Mytilidae. This family includes the most significant living coral borers. Bioerosion of living and dead corals is an important agent in the ecology of coral reefs. In particular, exploitation of living corals as a habitat was facilitated by the evolution of larvae to penetrate the living coral tissue, and thus of mechanisms to overcome the coral's defences and adult plasticity to keep pace with the growing coral (Morton 1990). Lithophaga sp. (Mytilidae, Lithophaginae) encloses a group of boring bivalves frequently recorded in both members of the Agrio Formation. Recent Lithophaga are chemical borers, a calcium carbonate chelating agent being produced from boring glands in the middle folds of each mantle lobe (Morton 1990). Boring occurs in thick-shelled bivalves, for example Myoconcha transatlantica Burckhardt, Gervillaria alatior (Imlay) and Isognomon (/.) ricordeanus (d'Orbigny), and also in coral colonies. Massive and ramose coral colonies usually have a high degree of infestation. Other boring activities on shells were detected through the whole Agrio Formation, but they remain to be studied in detail. Shallow-burrowing deposit-feeders. This guild is composed of irregular echinoids that ingest organic matter trapped in the substrate. Food resources include particulate organic detritus, living and dead smaller members of benthic flora and fauna, and organic-rich grains (Walker & Bambach 1974). Although locally abundant, irregular echinoids are restricted to a couple of horizons within the analysed section (Fig. 2). Nucleolites sp. (Cassiduloida, Nucleolitidae) is an irregular echinoid interpreted here as having a shallow infaunal deposit-feeding life habit. The lower density and smaller size of tubercles on the adapical side indicate mediumgrained sandy sediment (Smith 1984). It has simple phyllodes and lacks a masticator

apparatus (i.e. Aristotle's lantern), which is typical of regular echinoids. The phyllodes are a group of ambulacral pores located around the peristome. Their tube feet are responsible of the capture of nutritional particles and their transport to the mouth (Kier 1966). Nucleolites sp. is recorded in lithofacies 6 in the Agrio Formation at the Cerro Negro locality, central Neuquen. It occurs within the Holcoptychites neuquensis Subzone. Clypeopygus sp. (Cassiduloida, Nucleolitidae) is another irregular echinoid that shows clear features of a shallow infaunal deposit-feeding life habit. It has well-developed phyllodes that indicates the presence of specialized oral tube feet in collecting food particles from the sediment. It has also an incipient bourrelet, which is an externally inflated adoral part of the interambulacral areas around the peristome. This structure increased the surface cover with spines in that zone, helping to catch and transport the food particles towards the mouth. The reduced number of tubercles in the adapical side probably indicates that the zone of the petals would have been uncovered or with a very thin layer of sediment that allowed the passage of the tube feet through it. The periproct is supramarginal, posterior and located in a furrow, which drive the discharge off the respiratory tube feet (Smith 1984). The reduced size of the peristome indicates the absence of an Aristotle's lantern, which is absent in the great majority of the irregular echinoids. Clypeopygus sp. is recorded in the Agrio Formation in the Cerro Lotena locality, central Neuquen. It occurs within the Pseudofavrella angulatiformis and Olcostephanus laticosta zones. Pygorhynchus sp. (Cassiduloida, Nucleolitidae) is also interpreted here as a shallow infaunal deposit-feeder. It has well-developed petals with a slight tendency to close distally, well-developed, broad phyllodes and bourrelet, and a reduced pentagonal peristome located in anterior position. The periproct is marginal-inframarginal. Pygorhynchus sp. is recorded in lithofacies 6 in the Agrio Formation in Agua de la Mula within the Paraspiticeras groeberi Zone (Fig. 2). Deep-burrowing deposit-feeders. This guild is composed of anomuran decapod crustaceans, recorded by thalassinoids tentatively assigned to Protaxius sp. When only the chelipeds are present, placement in Protaxius sp. is problematic and part of the fossils may belong to Protocallianassa (Aguirre-Urreta 1989, 2003). Modern related genera of these burrowing shrimps have a complex pattern of behaviour. Callianassa builds galleries connected with each other that end at the sediment surface.

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They are mainly built in muddy sands or muds, up to 30 cm deep. The animal is a depositfeeder and the burrow serves as shelter, built in a way that the oxygen-rich waters can be pumped along it and make it permanently habitable. Callianassa uses the first chelipeds as the main digging tool. They loosen the sediment and move it backwards while the other walking limbs are used for scraping and pushing. Several behavioural states have been surveyed in Callianassa, including walking, carrying, dropping, stirring, grooming and turning (Stamhuis et al 1996, fig. 1). The diameter of the burrow exceeds the size of the body and the animal can move within it; at the branching points some sections are widened allowing the animal to turn in the burrow (Schafer 1972). Only occasionally fossil remnants of thalassinoid shrimps are found and these correspond to the first chelipeds, which are the only part of the carapace heavily calcified. They are most probably part of moultings. Callianassa moults inside the burrow and the tender exuviae is taken outside the burrow, except for the claws which are very heavy and rounded to be grasped and are left behind in the burrow (Schafer 1972). Callianassa, as many other burrow-dwelling animals, leaves the burrow with approaching death, and therefore corpses are rarely found inside the galleries and the heavy chelipeds are embedded outside the burrow if preserved at all. Protaxius sp. (and IProtocallianassa sp.) have been recovered from different stratigraphic levels in the Agrio Formation, especially in the Pilmatue Member. They are commonly preserved in small, near-spherical calcareous nodules with only the cheliped pair preserved; less often, just one of the chelipeds is preserved. The nodules are invariably embedded in shales of lithofacies 2. At few distinctive levels, especially in the Hoplitocrioceras gentilii Zone, specimens are complete and articulated, with the body parallel to the main axis of the nodule. These nodules, ovoid-irregular in shape, are preserved in fine sandstones of lithofacies 4 associated with burrowing systems of Thalassinoides. It is most probable that the decapod did indeed inhabit those galleries, comparable to the example documented by Bromley & Asgaard (1972) for the palinuran Glyphea rosenkrantzi Van Straelen. As stated above, this is an uncommon situation with thalassinoid shrimps, and the preservation of complete carapaces within the burrow system may indicate a sudden mass mortality. Nodules with complete shrimps are also found reworked and redeposited in coquinas of lithofacies 5.

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Semi-endobenthic Endobyssate suspension-feeders. This guild is composed of non-burrowing bivalves that lived partially buried in soft sediments and fixed by byssus threads. Three bivalve superfamilies are included in this guild (Fig. 3). Specific examples are Modiolus cf. subsimplex d'Orbigny, Pinna (P.) robinaldina d'Orbigny, Gervillella aviculoides (J. Sowerby) and Myoconcha transatlantica Burckhardt. P. robinaldina (Pinnoidea, Pinnidae) and M. cf. subsimplex (Mytiloidea, Mytilidae) are recorded as articulated shells or internal moulds in mudstones, and reworked in sandstones and coquinas. Clusters of P. robinaldina in life position occur on top of lithofacies 4. M. transatlantica (Carditoidea, Permophoridae) is locally abundant within lithofacies 5. It is frequently recorded articulated, bioeroded and out of life position. G. aviculoides (Pteriodea, Bakevelliidae) is locally abundant in lithofacies 5. It is a non-twisted monomyarian, and has an almost equivalve shell and an elongated outline. It occurs commonly out of life position as articulated fragments of the anterior region of the shell.

Epibenthic Free-lying suspension-feeders. This guild is composed of non-burrowing pteriomorphians that rest on, or are partially buried in, soft and firm substrata without attachment. Two superfamilies are included within this guild (Fig. 3). Aetostreon sp. (Ostreoidea, Gryphaeidae) encloses a group of thick-shelled oysters recorded in lithofacies 2 and 5 in the Pilmatue and Agua de la Mula members of the Agrio Formation. It has an inequivalve shell, with a very convex left valve and a slightly concave right one. The left valve is partly buried within the substrate, while the right valve has an opercular function. The lower valve is thicker than the upper one and has coarse commarginal lameleae. It corresponds to the cup-shaped recliners group. Inoceramus sp. (Pterioidea, Inoceramidae) encloses a group of thin- and flat-shelled inoceramids recorded in the lower portion of the Agua de la Mula Member of the Agrio Formation. It occurs in lithofacies 2 in pavements and within calcareous nodules, and is occasionally reworked in sandstones and coquinas. Mudstones bearing inoceramid bivalves have low benthic diversity, but they were probably deposited under normal oxygen levels as encrusted shells are recorded. Instead, very soft substrates are proposed as the main control that lowered the diversity (Wignall 1993). Inoceramid shells have an

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increased surface area to volume ratio interpreted as an adaptation to 'float' on very soft substrates. Mobile browsers. This guild is composed of prosobranch gastropods, and regular and irregular echinoids. Browsers are herbivores that scrape plant material from the depositional surface or chew or rasp (scrape) larger plants. Browsers commonly eat some superficial detritus too, so this category is actually transitional to deposit-feeding (Walker & Bambach 1974; Mangano & Buatois 1999). Pleurotomaria gerthi Weaver is a pleurotomariid gastropod (Prosobranchia, Archaeogastropoda) rarely occurring in the Pilmatue Member of the Agrio Formation. Shells are reworked and deposited in coquinas of lithofacies 2 and 5. They are usually encrusted on the external surface. Tylostoma jaworskii Weaver is an ampullospirid gastropod (Prosobranchia, Mesogastropoda) mostly recorded in the Pilmatue Member of the Agrio Formation. Shells are reworked and deposited in sandstones and coquinas. Observations on Cernina fluctuata (Sowerby), the sole extant species of the Mesozoic and Cenozoic Family Ampullospiridae, show that nocturnal in nature, these snails remain in burrows during daylight hours and emerge at dusk to browse over the substrate. Food resources include different macroalgae. Ampullospirids were previously assigned to drilling carnivorous Naticidae, but now are regarded as a separate but related family (Kase & Ishikawa 2003). Coenholectypus sp. (Holectypoida, Holectypidae) is an irregular echinoid with a large periproct located inframarginally. It shares some characters with the regular echinoids that makes the interpretation of its life habit and mode of feeding difficult, and there are no Recent relatives. It has a large peristome that probably carried a well-developed Aristotle's lantern, typical of epifaunal regular echinoids. However, the conical profile and the presence of tubercles that increase their size and number towards the oral side suggest a shallow infaunal life habit as in other irregular echinoids. Also, the ambulacral pore distribution around the peristome resembles simple phyllodes, suggesting that these echinoids could feed within the substrate (Zaghbib-Turki 1989). Smith (1984) interpreted the life habit of a similar and related genus as epifaunal forager during night hours and infaunal, relatively inert during daylight hours for protection from predators. Coenholectypus sp. is recorded from lithofacies 6 at Agua de la Mula. It occurs associated with Pygorhynchus sp. within the Paraspiticeras groeberi zones (Fig. 2).

Leptosalenia sp. (Calycina, Saleniidae) is the only regular echinoid recorded in the Agua de la Mula Member of the Agrio Formation and occurs slightly reworked in lithofacies 7 (Fig. 2). Although not preserved, a large peristome suggests that an Aristotle's lantern was present and that these echinoids were epifaunal scrapers. The presence of very large tubercles (one tubercle per plate) equally distributed on both sides of the skeleton also suggest an epifaunal mode of life. These tubercles supported large spines. On the oral side the spines were mainly responsible for the movement of the echinoid, but they were also used to abrade the substrate. Mobile carnivores. This guild includes mobile carnivorous decapod crustaceans that walk and crawl on the substrate and prey or scavenge on smaller organisms. They are represented both by palinurans and astacideans. The palinurans correspond to Meyerella rapax (Harbort), recognized in the Pilmatue Member of the Agrio Formation at several localities (Aguirre-Urreta 1989). M. rapax belongs to the Mecochiridae, a family of decapod crustaceans that became extinct in the Upper Cretaceous, within the Superfamily Glypheaoidea. At Cerro La Parva, specimens of M. rapax are highly abundant, being preserved in incomplete calcareous nodules that have been reworked and deposited at the base of a thin storm deposit, together with the ammonite Karakaschiceras attenuatus (Behrendsen) and oysters. The preservation closely resembles that recorded by Simpson & Middleton (1985) for M. magna M'Coy in the Atherfield Clay Formation of England. Specimens of M. rapax from Cerro La Parva are articulated, with the calcareous matrix not completely enclosing the carapace, thus producing a lateral compression of the fossil. Simpson & Middleton (1985) documented the presence of encrusting organisms on M. magna, such as the bivalves Anomia and Aetostreon, and suggested an epibenthic way of life for this lobster species. The material analysed here shows the encrusting oyster Amphidonte (Ceratostreon) sp., but it is difficult to determine whether the oysters settled on the specimens while alive, or attached to the already formed nodules. Preservational history of M. rapax probably included, following the death of the animals, an initial stage of shallow burial of corpses to produce the incomplete nodules, a relatively extended period of exposure of nodules on the seafloor to allow settlement and growth of the encrusters, and the reworking and final deposition of nodules at the base of a storm coquina. At

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other localities, incomplete specimens (mostly cephalothorax) are enclosed in flattish, ovoid, calcareous nodules. Most of the astacidean decapod crustaceans live today (and probably have lived in the past) in rock shelters and reefs, in high-energy littoral-sublittoral environments, but their fossil record is scarce. Although less abundant, the remains recorded usually represent the record of species that lived in low-energy environments in areas with quiet sedimentation (Forster 1985). The fossils are represented by nearly complete or undeformed individuals within calcareous nodules. They are represented by erymids and nephropids. The Erymidae are a group of extinct lobsters and Eryma has been characterized as a bottom-dwelling reptantian with four walking legs by Forster (1966). The first pereiopod is provided with claws and serves for protection and feeding. Many of them were probably also able to swim backwards as escape movements produced by curving the abdomen with its tail fan (Schafer 1972). Eryma sp. is only known from some horizons in the Pilmatue Member where the carapace of the cephalothorax is normally preserved, as well as isolated chelipeds, both in calcareous nodules and silty shales, but the abdomen is rarely found. Some of them are probably exuviae. The nephropids are much less common and are only represented by the homarinae Hoploparia sp. The close living relative is Homarus, the common lobster, with a few species. They are carnivorous and may cover wide distances to find dead or living prey, occasionally breaking open thick-shelled gastropods. They live generally on rocky grounds, but also occur in sandy and muddy areas of the open seafloor (Schafer 1972). Hoploparia sp. is recorded in the Pilmatue Member and the basal beds of the Agua de la Mula Member. All the fossils correspond to isolated (very rarely paired) chelipeds preserved in calcareous nodules within shales of lithofacies 2. Epibyssate suspension-feeders. This guild is represented by non-burrowing pteriomorphian bivalves whose adults lived attached by a byssus and the foot has usually lost its function. Only those species that spend most of the time fixed are included here and stability is achieved through the byssal fixation. Five superfamilies are included in this guild (Fig. 3). Specific examples are Mytilus sp., Gervillaria alatior (Imlay), Isognomon (/.) ricordeanus (d'Orbigny), /. (I.) lotenoensis (Weaver) and Mimachlamys robinaldina (d'Orbigny). G. alatior (Pterioidea, Bakevelliidae) is a locally abundant monomyarian species preserved

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commonly as articulated shells. It has a very thick shell, highly inequivalved (with the left valve more convex than the right), and slightly twisted. It has an antero-dorsal triangular groove separating the dentition from the anterior margin of the shell (on internal shell surface) interpreted as a byssal groove. Although byssate, the main stabilizing factor in the adult stage seems to be the heavy umbo, leaving the bivalve completely sessile. M. robinaldina (Pectinoidea, Pectinidae) is the most common pectinid species. It occurs throughout the sequence. Articulated shells are rare, while dissarticulated valves with fragmented auricles are common and the species is one of the main components of shell debris. It has a slightly inequivalved shell (left valve more convex than right), asymmetric auricles on right valve, and right anterior byssal notch. Limid bivalves (Limoidea) are scarcely recorded in the Agrio Formation. They have an equiconvex shell and small auricles without byssal notch. They are recorded as single valves with broken auricles. Cemented suspension-feeders. This group includes cemented serpulids, bivalves and bryozoans, which may have solitary, colonial or gregarious life habits. Bivalves are represented by pteriomorphians. Two bivalve superfamilies are involved (Fig. 3). Amphidonte (Ceratostreon) sp. (Ostreoidea, Gryphaeidae) encloses a group of small- to medium-sized cemented oysters with radially ribbed shells. Small tubercles and pits (chomata) are present on the inner margin of both valves. They lived cemented by the left valve. It is the most common bivalve of the Agrio Formation, and it is found encrusting ammonoids, nautiloids, corals, serpulids and bivalves. In addition, shell debris is commonly oyster-dominated. Its life habit may be gregarious as different specimens cement to each other forming banks similar to recent oyster reefs. Plicatula sp. (Pectinoidea, Plicatulidae) is rare in the Agrio Formation, attached to the external surfaces of bivalve shells. Encrusting bryozoans (Cyclostomata, Tubuliporina) are also part of this guild. The colonies formed a thin cover over their substrate and attached directly to it. They are recorded encrusting bivalve shells. Serpulids (Annelida, Polychaeta, Sabellida) are worms that build a calcareous habitation tube, which is generally attached to a submerged surface. Most living taxa are marine, with only one living species known from freshwater (Ten Hove & van den Hurk 1993). Parsimonia antiquata (J. de C. Sowerby) is the most common serpulid tube of the Agrio Formation. It is

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recorded as solitary tubes or in small gregarious clusters. Cementation occurs on valves, shells and calcareous nodules. Tubes are closed at one end and increase in diameter to a maximum of 13 mm. They are straight-weakly sinuous, but in some circumstances they can be coiled. Within the 'Neocomites' sp. Subzone there are shell beds embedded in lithofacies 2 composed almost entirely of gregarious P. antiquata. Well-preserved tubes and shells suggest a low degree of lateral transport. These shell beds have probably a mixed sedimentary-biological origin under particular palaeoenvironmental conditions, and their stratigraphic and palaeoecological significance (i.e. sedimentation rate, turbidity, nature of the bottom, nutrients availability, temperature and salinity) remain to be analysed in detail. Sarcinella occidentalis (Leanza & Castellaro) is a serpulid with colonial life habit, characterized by bundles of tubes generated by asexual reproduction. The association of bundles is fan-shaped (70 cm long, 55 cm width and 15 cm thick). Each bundle is composed of approximately 30 near-parallel tubes, up to 1 mm in diameter. Bundles are cylindrical and have a ramose aspect. S. occidentalis is rare in the Agrio Formation and far less common than P. antiquata. Isolated and fragmented centimetre-thick bundles occur in lithofacies 5 and associated bundles are very occasionally recorded. Cemented or rooted suspension-feeders. This guild is composed of erect bryozoans and Demospongea, both recorded in the coral patch reef facies of the Agua de la Mula Member. Sponges and bryozoans are accessory elements within the reef, but the latter can be dominant in intercalated coquina beds. Attachment to substrate could be either cemented or rooted by small firm tubes or hair-like radicles. Erect bryozoans, ramose in aspect, have zooecial apertures from the base to the tip of the colony. Colonies are a few centimetres long and 5 mm wide. Erect growth strategy minimizes competition for substrate and provides access to nutrients higher in the water column (Hageman et al. 1998). Sponges are tubular-shaped, 2 cm in height and 1 cm in width, and have an osculum in the upper surface. Cemented microcarnivores. This guild is composed of scleractinian corals that consist mainly of colonial species. A coral patch reef is recorded at the top of the Agua de la Mula Member within the Paraspiticeras groeberi Zone (Fig. 2, see lithofacies 7). Coral colonies are also abundant in the Holcoptychites agrioensis Subzone of the

Pilmatue Member. They consist of reworked, small globose and branching colonies, and occur in lithofacies 5. Apart from these records, corals are rare in the Agrio Formation.

Nekto-benthic Scavengers. This guild is represented by Cymatoceras perstriatum (Steuer), the only nautiloid recorded in the Agrio Formation (Cichowolski 2003). It is interpreted here that these animals were slow swimmers that lived associated with the bottom, on which they fed mainly scavenging and where egg-lying occurred. C. perstriatum is not associated to reef facies, and its depth habitat would be much less than that of the recent Nautilus. The septal strength index (SSI) for this species correlates with depths of about 350 m. This index assesses the implosion depth of different morphologies based on the septal thickness and curvature radius (Westermann 1973). Therefore, it seems probable that C. perstriatum lived in a rather different environment to that of their living relatives (see Saunders & Ward 1987). Nautiloids occur in the Agrio Formation in lithofacies 2 (in the intercalated sandstones and coquinas), 4 and 5. They have never been found in shales of lithofacies 1 and 2, while the ammonoids are recorded frequently in shales. The preservation of cephalopods in shales is influenced by the post-mortem capacity to float of shells, which is related to morphological traits such as the length of the body chamber, siphuncle width and position and number of septa per whorl. Ammonoid shells probably sunk after death faster than nautiloid shells, which floated for a longer period of time (Reyment 1958), thus potentially ending up in shallower environments by the action of waves and currents.

Nektonic Carnivores This guild is composed of ammonoids, pycnodontiform fishes, elasmosaurid plesiosaurs and ichthyosaurs. They were carnivores that swam within the water column looking for a prey to hunt. Ammonoids were probably carnivores (Monks & Palmer 2002), but not pursuit predators as modern tuna or squids. They may have been ambush predators: they lie passively waiting for prey and then accelerate rapidly to grab them as modern cuttlefishes (Jacobs & Chamberlain 1996; Monks & Palmer 2002). The relatively large size of the buccal mass of ammonoids

GUILDS IN THE LOWER CRETACEOUS

suggests that they ingested whole small organisms or bite pieces of larger prey. Remains identified in situ in crop and stomach contents in body chambers of ammonites include fragments of crinoids, cephalopod beaks and ostracods (Nixon 1988, 1996). Heteromorphic forms must have been poorer swimmers than planispiral forms, and among the latter compressed involute shapes were good for continuous swimming in comparison with less compressed shells (Chamberlain 1981). Newly hatched ammonoids may have been swimmers or passive vertical migrators in the plankton, drifting with surface currents (Landman et al. 1996). Several data suggest that these very young ammonoids may have lived in a different environment from that of older juveniles and adults. Ammonoids are present in both marine members of the Agrio Formation. They are most common in shales of lithofacies 2, although they are also recorded in lithofacies 1, 3, 4 and 5. The ammonoid fauna in any given horizon is generally represented by few species, some are even monospecific. The different species are normally present in more than one facies, and their mode of preservation varies accordingly. The best faunas include complete specimens with shells, embedded in dark shales such as Olcostephanus (O.) atherstoni (Sharpe), Karakaschiceras spp. or encased in calcareous nodules, for example Olcostephanus (O.) laticosta (Gerth) and Hoplitocrioceras gentilii Giovine. Sometimes the nodules are heavily encrusted with oysters and serpulids, suggesting a long period of residence on the seafloor. There are also levels where only body chambers are preserved, for example Chacantuceras ornatum Aguirre-Urreta & Rawson and Spitidiscus riccardii Leanza & Wiedmann. The preservation is rather poor in sandstones and coquinas, where shells are usually encrusted by oysters and serpulids. Typical examples are Holcoptychites agrioensis (Weaver) and Crioceratites diamantensis / C. andinus (Gerth). The Pycnodontiformes (Osteichthyes, Actinopterygii, Neopterygii) were a chiefly Mesozoic, sister group to Teleostei. They were mostly deep bodied, laterally compressed fishes with pectoral fins located up on the sides (Nursall 1999). They have a specialized dentition composed of a series of crushing teeth arranged in longitudinal rows on the median vomer in the roof of the mouth, and similar rows on the paired prearticulars of the lower jaw (Nursall 1999). They probably fed on molluscs, echinoids and other hard-bodied preys. Most pycnodontiform remains were found in exclusively marine deposits. Bocchino (1977) described and

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figured Macromesodon agrioensis nov. sp. based on a pycnodontiform vomerian dentition recorded in the Pilmatue Member of the Agrio Formation. Another nearly complete dentition and isolated teeth have been found in lithofacies 5 in the Holcoptychites agrioensis Subzone. The first Lower Cretaceous plesiosaur remains from Argentina came from the Pilmatue Member (Lazo & Cichowolski 2003). The remains consist mainly of vertebral material and other fragmented bones of the post-cranial skeleton. They show clear affinities to the Family Elasmosauridae. They were found disarticulated at four different levels ranging in age from the Late Valanginian to the Early Hauterivian (Figs 1 & 2). Three of the levels belong to lithofacies 2, the other one to lithofacies 5. In addition, few isolated vertebrae of ichthyosaurs have been found in the Pilmatue Member (Cichowolski & Lazo 2000). Summary of the depositional history of the Agrio Formation It is important to note that the interpretation of the palaeoecology and palaeoenvironments of the Agrio Formation discussed here is restricted to the Agua de la Mula locality and adjacent areas, and is based on the analysis of sedimentary facies and macrofossils. This discussion is also constrained to a generalized interpretation based on the analysis of more than 1200 m of sequence. The details of particular variations in the facies and fossil assemblages are beyond the scope of this paper. In the study area the Agrio Formation was deposited in a ramp setting ranging from lowenergy basin to high-energy inner ramp, but its subenvironments are not equally represented. The outer (lithofacies 2) and mid-ramp (lithofacies 3, 4 and 5) are better represented, and they record frequent storm reworking and deposition. They developed under normal marine conditions, as indicated by lithofacies, macrofossils, and normal oxygen level at the seafloor inferred from the diverse and abundant benthic macrofauna. The substrate consistency varied between soft and firm due to periods of low net rate of sedimentation by omission (bypassing or starvation) or erosion (winnowing or reworking). Islands of hardgrounds are occasionally recorded as the result of prolonged omission, cementation, erosion, accumulation of shell gravels or exposure of calcareous nodules on the seafloor. Water temperatures were probably warm, as indicated by the presence of thick-shelled bivalves, oolites and coral patch reef. The basin

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(lithofacies 1), inner ramp (lithofacies 6) and coral patch reef (lithofacies 7) are of minor significance, and are restricted to the Agua de la Mula Member. The basin was the only sector of the ramp in which low oxygen levels existed at the seafloor, while on the inner ramp deposition took place in a very shallow, high-energy sea during times of reduced siliciclastic input.

The trophic structure of the benthic invertebrates is clearly dominated by suspension-feeders (nine guilds) reflecting the availability of suspended food particles. Suspension-feeders formed a tiering structure ranging from deep levels in the sediment to well above the sediment-water interface. Deposit-feeders and browsers are less abundant, being restricted to two infaunal

Fig. 4. Reconstruction of a fossil assemblage recorded in shales (lithofacies 2) in the P. angulatiformis Subzone in the Pilmatue Member of the Agrio Formation. The faunal composition is based on field observations, and the life habits are the result of the study of the functional morphology of skeletons, on taphonomic observations and on comparisons with Recent relatives, a, Amphidonte (Ceratostreori) sp.; b, Cucullaea (Noramya) gabrielis Deshayes in Leymerie; c, Disparilia sp.; d, Steinmanella pehuenmapuensis Leanza; e, Veneridae sp. indet.; f, Protohemichenopus neuquensis Camacho; g, Pseudofavrella angulatiformis (Behrendsen); and h, Parsimonia antiquata (J. de C. Sowerby).

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Fig. 5. Reconstruction of the fossil assemblage recorded in sandstones of lithofacies 3 in the Holcoptychites agrioensis Subzone in the Pilmatue Member of the Agrio Formation. The faunal composition is based on field observations, and the life habits are the result of the study of the functional morphology of skeletons, on taphonomic observations and on comparisons with Recent relatives, a, Ramose and massive coral colonies; b, Cucullaea (Noramya) gabrielis Deshayes in Leymerie; c, Gervillaria alatior (Imlay); d, Isognomon; (/.) ricordeanus (d'Orbigny); e, Mimachlamys robinaldina (d'Orbigny); f, Myoconcha transatlantica Burckhardt; g, Pinna (P.) robinaldina d'Orbigny; h, Pterotrigonia coihuicoensis (Weaver); i, Ptychomya koeneni Behrendsen; j, Steinmanella transitoria (Steinmann); k, Trigonia carinata Agassiz; 1, Cymatoceras perstriatum (Steuer); m, Holcoptychites agrioensis (Weaver); n, thalassinoid decapod crustacean; and o, pycnodontiform fish.

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guilds and one epifaunal guild, respectively. Based on field observations, three different fossil assemblages are reconstructed in Figures 4-6 to illustrate how the guilds associate with the lithofacies (see also Lazo 2004). A fossil assemblage of the outer ramp is illustrated in Figure 4. It comes from the Pseudofavrella angulatiformis Subzone in the Pilmatue Member. The outer-ramp assemblages were characterized by a relatively low benthic

diversity. The benthic macrofauna is dominated by shallow burrowers and islands of cemented epifaunal organisms. This fauna inhabited a rather quiet and well-oxygenated muddy bottom, which was occasionally altered by distal storm processes. Nektonic ammonoids and marine reptiles complement the association. A fossil assemblage of the mid-ramp is illustrated in Figure 5. It occurs in the Holcoptychites agrioensis Subzone in the Pilmatue Member.

Fig. 6. Reconstruction of the coral patch reef lithofacies recorded in the Paraspiticeras groeberi Zone in the Agua de la Mula Member of the Agrio Formation. The faunal composition is based on field observations, and the life habits are the result of the study of the functional morphology of skeletons, on taphonomic observations and on comparisons with Recent relatives, a, Demospongea; b, ramose corals; c, globose corals; d, flat corals; e, conical corals; f, erect bryozoans; g, Amphidonte (Ceratostreori) sp.; h, Lima sp.; i, Leptosalenia sp.; and j, Gastrochaenolites isp.

GUILDS IN THE LOWER CRETACEOUS Mid-ramp assemblages were frequently reworked by storm processes and deposited in tempestites. They are characterized by a high diversity, and a well-developed tiering structure that includes deep and shallow infaunal, semi-infaunal, epifaunal, nekto-benthic and nektonic organisms. Coral colonies are abundant only in one horizon. They are small in size and grew over shell gravels, with a patchy distribution. They represent an incipient colonization stage in unfavourable conditions for reef growth. Swimming organisms are dominated by cephalopods, while fish and marine reptiles are rare. The coral patch reef association of the upper midramp, recorded in the Agua de la Mula Member, is reconstructed in Figure 6. It is dominated by epibenthic organisms, mostly cemented, rooted or byssate, but there are also mobile animals such as regular echinoids. The corals built a positive topographic structure, which was probably capable to resist hydrodynamic stress. This association represents particular palaeoceanographic conditions allowing the growth of the reef.

Conclusions •

• •

•

•

The guild analysis indicates normal benthic oxygen level and overall normal salinity. Benthic suspension-feeders, are more common than deposit-feeders, suggesting the predominance of suspended food particles over deposited food resources. Trace fossils indicate soft-firm muddy and sandy bottoms. Most of the fossil associations were autochthonous or slightly parautochthonous, which in turn means a low degree of lateral transport by storms. This is important to infer different palaeoenvironmental conditions based on fossil composition. This situation could be expected in other storminfluenced settings. The observed taphonomic pattern points to a within-habitat time-averaging of fossils. This situation could also be expected in other storm-influenced settings. The study setting was mainly marine and almost permanently connected to the Pacific Ocean, but periodic subtle changes in the siliciclastic (or freshwater) input and changes in salinity did, indeed, occur. These changes are probably related to relative sea-level changes and enhaced the development of unfrequent fossil associations, such as coral patch reefs and serpulid aggregations.

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The authors thank G. lovine, S. Lanes, M. Lantz, P. Rawson, C. Rodriguez and R. Sanci for their help in the field. R. Palma is acknowledged for his help in the description and interpretation of thin sections, and A. Zamorano for drawing the fossils. The manuscript

has been improved by the comments and suggestions of F. Fursich, P. Wignall, G. Veiga and J. A. Howell. This study was partially funded by UBACYT X 084 and ANPCYT 14143/03 grants to M.B. Aguirre-Urreta and

IAS postgraduate grant to D. Lazo.

References AGUIRRE-URRETA, M.B. 1989. The Cretaceous decapod Crustacea of Argentina and the Antarctic Peninsula. Palaeontology, 32, 499-552. AGUIRRE-URRETA, M.B. 2003. Early Cretaceous decapod Crustacea from the Neuquen Basin, west-central Argentina. Contributions to Zoology, 72, 79-81. AGUIRRE-URRETA, M.B. & RAWSON, P.P. 1997. Th ammonite sequence in the Agrio Formation (Lower Cretaceous), Neuquen Basin, Argentina. Geological Magazine, 134, 449-458. AGUIRRE-URRETA, M.B., RAWSON, P.P., CONCHEYRO, G.A., BOWN, P.R. & OTTONE, E.G. 2005. Lowe Cretaceous (Berriasian-Aptian) biostratigraphy of the Neuquen basin. In: VEIGA, G.D., SPALLETTI, L.A., HOWELL, J.A. & SCHWARZ, E. (eds) The Neuquen Basin: A Case Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 57-81. AIGNER, T. 1982. Calcareous Tempestites: Stormdominated stratification in Upper Muschelkalk Limestones (Middle Trias, SW-Germany). In: EINSELE, G. & SEILACHER, A. (eds) Cyclic and Event Stratification. Springer, Berlin, 180-198. BOCCHINO, A. 1977. Un nuevo Gyrodontidae (Pisces, Holostei, Pycnodontiformes) de la Formacion Agrio (Cretacico Inferior) de la Provincia de Neuquen, Argentina. Ameghiniana, 14, 175-185. BRINKMANN, H.-D. 1994. Facies and sequences of the Agrio Formation (Lower Cretaceous) in the central and southern Neuquen Basin, Argentina. Zentralblatt fur Geologic und Palaeontologie, I, (1/2), 309-317. BROMLEY, R.G. & ASGAARD, U. 1972. The burrow and microcoprolites of Glyphea rosenkrantzi, a Lower Jurassic palinuran crustacean from Jameson Land, East Greenland. Gronlands Geologiske Unders0gelse, Rapport, 49, 15-21. CHAMBERLAIN, J.A., JR. 1981. Hydromechanical design of fossil cephalopods. In: HOUSE, M.R. & SENIOR, J.R. (eds) The Ammonoidea. Systematics Association, Special Volume, 18, 289-336. CICHOWOLSKI, M. 2003. The nautiloid genus Cymatoceras from the Cretaceous of the Neuquen and Austral basins, Argentina. Cretaceous Research, 24, 375-390. CICHOWOLSKI, M. & LAZO, D.G. 2000. Lower Cretaceous marine reptiles from Argentina. In: Abstracts 31st International Geological Congress, Rio de Janeiro, Brazil, August 6-17. On CD.

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Mid-Cretaceous turnover of saurischian dinosaur communities: evidence from the Neuquen Basin RODOLFO A. CORIA1 & LEONARDO SALG 1

CONICET, Secretaria de Estado de Cultura de Neuquen, Museo Carmen Funes, Av. Cordoba 55, (8318) Plaza Huincul, Neuquen, Argentina (e-mail: [email protected]) CONICET, Museo Universidad Nacional del Comahue, Buenos Aires 1400,,, (8300) Neuquen, Argentina (e-mail: [email protected]) Abstract: Successional changes are recorded among saurischian dinosaur faunas throughout the Cretaceous of the Neuquen Basin. Both sauropod and theropod lineages present in NW Patagonia show comparable transformations and indicate that some clades did not survive the Late Turonian boundary. Similar patterns of changes are observed in other Patagonian sedimentary basins. Diplodocoid sauropods and carcharodontosaurid theropods are not yet recorded in levels younger than Turonian, while titanosaurians and abelisauroids are present before and after this period. Important floral changes during the mid-Cretaceous, mainly related to the proliferation of angiosperms, could be one element that influenced some dinosaurian associations.

The Neuquen Basin is perhaps one of the most thoroughly prospected areas in Argentina, not only for natural resources, but also for Mesozoic fossil vertebrates, especially dinosaurs. A great many South American dinosaur records are derived from strata in this basin. The first reference of dinosaurs in South America is of fossils collected from Late Cretaceous rocks of the Neuquen Group (Coria & Salgado 2000), the uppermost terrestrial sediments exposed in the Neuquen Basin. Since then, over 40 different dinosaur species (including Cretaceous birds) have been reported (Novas 1997; Coria 1999). This extensive sample of dinosaurian evolution, with a broad diversity and time span, is one of the few glimpses of evolutionary patterns from a single region that are probably applicable to a wider geographical range. The faunal changes revealed by dinosaur associations have intrigued many researchers but have been treated lightly in the literature (Novas 1997). One conspicuous example of these changes occurred during the transition from the Early to the Late Cretaceous. Taking into account a broad overview of the fossil evidence recorded in other sedimentary basins, it seems that this transition of dinosaurian faunas is documented throughout Patagonia as a whole. In this contribution, we describe the composition of the assemblages of saurischian dinosaurs, mainly immediately below and above the Turonian-Coniacian boundary, taking into

account that evolutionary changes in the flora could be a possible biological explanation for this record.

Sauropoda The oldest record of a sauropod in the Neuquen Basin is a partial hindlimb (femur, tibia and fibula) found in the Upper Jurassic (Kimmeridgian-Tithonian) Tordillo Formation of northern Neuquen Province (Garcia et al. 2003). The proximal end of the tibia is laterally compressed, contrasting the condition observed in neosauropods where the proximal end of the tibia is anteroposteriorly compressed (Wilson & Sereno 1998). This fossil suggests that basal eusauropods lived in northern Patagonia during Jurassic times, as has been already documented in the central Patagonian San Jorge Basin (Garcia et al 2003). Neocomian Patagonian sauropods are scarce, and the few samples available come from the Hauterivian La Amarga Formation of the western Neuquen Basin. The diplodocoid Amargasaurus cazaui (Fig. 1) is the oldest Cretaceous sauropod species recorded in the Neuquen Basin (Salgado & Bonaparte 1991). It is a member of the Dicraeosauridae, whose earliest representatives come from the Upper Jurassic of Africa. By the upper Lower Cretaceous, the sauropod record is characterized by the coexistence of

From: VEIGA, G. D., SPALLETTI, L. A., HOWELL, J. A. & SCHWARZ, E. (eds) 2005. The Neuquen Basin, Argentina: A Cas Study in Sequence Stratigraphy and Basin Dynamics. Geological Society, London, Special Publications, 252, 317-327. 0305-8719/057$ 15.00 © The Geological Society of London 2005.

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Fig. 1. Skeletal reconstruction of a typical diplodocoid sauropod from Patagonia, Amargasaurus.

titanosaurian-related forms and rebbachisaurids, a clade of basal diplodocoids widely distributed throughout the world (Salgado 2003a). This sauropod assemblage comes from several Lower and mid-Cretaceous units, such as the Lohan Cura and Candeleros formations (Salgado et al. 2004) (Fig. 2a-c). The oldest undoubted titanosaurian recorded in the Neuquen Basin, Andesaurus delgadoi (Fig. 3), is known from the latter (Calvo & Bonaparte 1991). In Patagonia, the oldest record of a titanosauroid (titanosaurian with procoelous caudal vertebrae: Salgado 2003/?) is Cenomanian in age (Calvo & Salgado 1998). The period between the upper Cenomanian and the latest Cretaceous is characterized by the dominance of eutitanosaurians (defined as all titanosaurs closer to Saltasaurus than to Epachthosaurus: Salgado 2003Z?). Saltasaurids (sensu Wilson 2002) are restricted to the Campanian-Maastrichtian. Pellegrinisaurus, Laplatasaurus and Neuquensaurus from the Rio Colorado Subgroup at Cinco Saltos are perhaps members of the Saltasaurinae (sensu Sereno 1998); that is, the clade comprising all saltasaurids closer to Saltasaurus than to Opisthocoelicaudia. Other saltasaurines such as Rocasaurus and Aeolosaurus are recorded in the Allen Formation. The placement of the second genus within the Saltasaurinae results from applying the Sereno (1998) definition of Saltasaurinae on the Salgado et al. (1997) cladogram. The Saltasaurini (= Saltasaurinae sensu Salgado et al. 1997), defined as Neuquensaurus, Saltasaurus, their most recent common ancestor and all of its descendants, are restricted, both spatially and chronologically, to the uppermost Cretaceous (Campanian-Maastrichtian) of southern South America. They are recorded in the Argentine provinces of Rio Negro, Neuquen

and Salta. The observed absence of Saltasaurini south of 42°S might be due to the imperfection of the palaeontological record, but it is also possible that the North Patagonian Massif impeded the passage of the Saltasaurini and prevented representatives of this clade from colonizing more southerly regions (Salgado & Azpilicueta 2000).

Theropoda Meat-eating dinosaurs were always less abundant than herbivores, at least in terms of biomass. However, the theropod fauna recorded from the Neuquen Basin is as diverse as the sauropod record. Even though we cannot be sure whether this numerical similarity is due to preservational biases or to a relatively higher taxonomic diversity of meat-eating dinosaurs, the theropod record provides an excellent example of successional change. The theropod record from the Neuquen Basin involves a number of taxa that may be clustered into three major clades: Abelisauroidea, Carcharodontosauridae and Coelurosauria. The greater diversity corresponds to abelisauroid forms. The best preserved South American abelisauroid is Carnotaurus sastrei (Fig. 4). Some of them have been recorded as old as Lower Barremian (Bonaparte 1996; Leanza etal. 2004). Abelisaur oids (sensu the definition given by Wilson et al. 2003) are basically a Gondwanan lineage of meat-eaters, whose highest diversity has been recorded in Patagonia, particularly in the different terrestrial Cretaceous horizons of the Neuquen Basin. The reports of Laurasian abelisauroids (Buffetaut et al. 1988; Acarie et al. 1995) still need to be confirmed on the basis of more complete specimens. Among the European

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Fig. 2. (a) Pubis, (b) femur and (c) mid-caudal vertebrae of Limaysaurus sp. (Lohan Cura Formation, Lower Cretaceous, Neuquen, after Salgado et al 2004). (d) Carcharodontosaurid tooth (MCF-PVPH-108) from the Huincul Formation, Early Late Cretaceous and (e) probably abelisauroid teeth MCF-PVPH-421 from Anacleto Formation, Late Cretaceous. Scales are in cm.

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Fig. 3. Skeletal reconstruction of a basal titanosaurian sauropod, Andesaums.

forms, Genusaurus could be related with more basal abelisauroids, but it is equally parsimonious to consider a sibling relationship with this taxon. In South America, the Abelisauroidea show a continuous presence throughout the Neuquenian section from the Lower Cretaceous up to the Campanian (Paulina Carabajal et al. 2003). The oldest abelisauroid record occurs in the Lower Barremian La Amarga Formation (Leanza et al. 2004) with Ligabueino andesi (Bonaparte 1996). Within the mid-Cretaceous units of the Neuquen Basin, besides diagnostic but fragmentary specimens (MCF-PVPH-237 and MCFPVPH-380, Museo Carmen Funes, Paleontologia de Vertebrados, Plaza Huincul, Neuquen, Argentina) from the Cerro Lisandro and Portezuelo formations, only Ilokelesia aguadagrandensis from the presumably Turonian Huincul Formation (Leanza et al. 2004) has been properly described

(Coria & Salgado 1998). This formation is probably correlated with the Lower Member of the Bajo Barreal Formation of the central Patagonian San Jorge Basin, which yields fragmentary but identifiable abelisauroid remains (Lamanna et al. 2002). These include the fragmentary Xenotarsosaurus bonapartei (Martinez et al. 1986) that has been provisionally regarded as a basal abelisauroid (Coria & Salgado 1998). In turn, in higher Cretaceous levels, the best preserved abelisauroid specimens collected correspond to the Abelisauridae. This is a group of large, lightly-built and highly derived theropods with occurrences in South America, Africa, Madagascar and India (Bonaparte 1991; Sampson et al. 1998; Wilson et al. 2003; Sereno et al. 2004). Up to the present, South America has provided the most diverse samples of abelisaurids, all of them from Campanian units. Whereas

Fig. 4. Skeletal reconstruction of the better-known abelisauroid theropod, Carnotaurus (modified from Bonaparte et al. 1990).

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Abelisaurus comahuensis and Aucasaurus garridoi are from the Campanian Anacleto Formation (Bonaparte & Novas 1985; Coria et al 2002), Carnotaurus sastrei (initially considered to provene from the Aptian Gorro Frigio Formation) is now regarded to be from the Campanian-Maastrichtian La Colonia Formation (Ardolino & Delpino 1987). The second main group of South American theropods is the Carcharodontosauridae, a group of theropod dinosaurs characterized by their large sizes and robust skeletons. These are represented by fewer genera than the abelisauroids, and only three reports may be considered as definitive samples of this taxon. Despite the fact that their remains have been claimed at several Patagonian localities (Alcober et al. 1998; Novas et al 1999; Veralli & Calvo 2003), the only Patagonian carcharodontosaurid formally named and preliminarily described so far is Giganotosaurus carolinii (Coria & Salgado 1995) (Fig. 5). Additional carcharodontosaurid specimens (Fig. 1) composed of partially preserved skeletons have been found within and outside the Neuquen Basin (Coria & Currie 1997; Rich et al. 1998; VickersRich et al. 1999), although their description awaits publication. Recently, carcharodontosaurids have been reported from Cenomanian-Turonian units, including teeth from the Portezuelo and Mata Amarilla formations (Novas et al. 1999; Veralli & Calvo 2003), and a partial skeleton from

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Late Cretaceous exposures of Mendoza Province (Rio Colorado Formation: sensu Alcober et al. 1998). Because these specimens have been not properly described and figured, precise evaluation and comparison with other material is not possible. The identification of the teeth is based on the presence of deep wrinkles on the labial sides of the teeth adjacent to the carinae (Fig. 2d). This feature has been proposed as a unique feature for Carcharodontosauridae (Sereno et al. 1996). The identification of the Mendoza specimen as a carcharodontosaurid (Alcober et al. 1998) was based on the presence of extreme axial pneumaticity, coosification of a number of medial gastralia as V-shaped struts and a hypertrophied pubic foot. Much of this information is field observation and most of the specimen remains unprepared. Moreover, virtually all the features mentioned match with a coelurosaur specimen recently found in levels related to the Portezuelo Formation at Sierra Barrosa, Neuquen Province (Coria & Currie 2002). The stratigraphy of the Cretaceous beds of the northern Neuquen Basin has not been studied in detail. Therefore, it is likely that the Mendoza theropod comes from levels older than the Rio Colorado Subgroup, mainly based on the great amount of shared osteological features between this form and the one described from Sierra Barrosa, which is definitively not a carcharodontosaurid. However, some caveats regarding the phylogenetic relevance of certain dental features

Fig. 5. Skeletal reconstruction of the Patagonian carcharodontosaurid Giganotosaurus.

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need to be taken into consideration. Personal observations of many theropod teeth suggest that tooth morphology may be linked with prey acquisition and processing, and that similar features developed independently in multiple theropod lineages (Fig. 2e). Wrinkles of variable degree of development can be observed in the tooth row of a single individual or even to be absent in some dental pieces (i.e. Giganotosaurus Holotype specimen, MUNC-CH-1). Therefore, more complete specimens need to be found to confirm the carcharodontosaurid record in several Patagonian localities. The third group of theropods recorded in the Neuquen Basin includes several coelurosaurs that seem to have more cosmopolitan distributions with relatives in South America, North America and eastern Asia (Novas 1997, 1998; Novas & Puerta 1997; Coria & Currie 2002; Chiappe & Coria 2003). So far, the record of these specimens is mostly restricted to Turonian-Coniacian units (i.e. Portezuelo Formation: Leanza et al 2004). The presence of flightless theropods on continents that were geographically isolated during the Late Cretaceous has yet to be explained. However, in this contribution we have focused on the abelisauroid and carcharodontosaurid faunal dynamics, and their similarity with the turnover of diplodocoid and titanosaurian faunas.

Discussion In Europe (Le Loeuff 1991), Africa (Jacobs et al 1993; Sereno et al. 1998, 1999) and North America (Lucas & Hunt 1989), the mid-Cretaceous (Aptian-Cenomanian) seems to coincide with an apparent extinction of some saurischian lineages: basal diplodocoids and basal titanosaurians, carcharodontosaurids and spinosaurs on the first two continental land masses; and basal titanosauriformes (Tidwell et al. 2001) and carcharodontosaurids (if Acrocanthosaurus is corroborated as a carcharodontosaurid, as proposed by Sereno et al. 1996, but see Currie & Carpenter 2000 and Coria & Currie 2006) on the third. Although evidence is still insufficient, it is possible to hypothesize that a single series of events might be responsible for the extinction of these European, African and North American saurischians, and the near-simultaneous disappearance of diplodocoids, basal titanosaurians and carcharodontosaurid theropods from the Patagonian fossil record. However, although the dinosaur turnover was probably of global scale, the groups that proliferated after the midCretaceous extinction were different in each

case (Bakker 1978; Britt & Stadman 1997; Novas et al. 2005). Reasonably, we can assume that these hypothetical events involved environmental and floral changes, and that the immediate effect on the sauropods (herbivores) affected the populations of theropod (meat-eating) dinosaurs. Unfortunately, the specific causes of the faunal changes recorded in the mid-Cretaceous, as well as the nature of the environmental modifications, are far from being comprehensible, in part because of the lack of data, which is especially true for the Neuquen Basin. In spite of this weighty restriction, some authors have ventured partial explanations for the faunal turnover involving dinosaurs recorded elsewhere by the mid-Cretaceous. For example, Lucas & Hunt (1989) have linked the apparent extinction of North American Lower Cretaceous sauropods (mostly basal titanosauriforms) to climatic changes produced by a Late Albian regression. In fact, extinction of non-marine and marine faunas at the Cenomanian-Turonian boundary has been recorded in SW Utah (Eaton et al. 1997), and, according to Lucas & Hunt (1989), it is possibly connected with the extinction of North American sauropods. In Africa, the maximum transgression was near the end of the Cenomanian, just before the definitive break between this continent and South America, and coincident with the major rise in sea level worldwide (Benton et al. 2000). Importantly, the saurischians recorded in the Cenomanian of northern Africa (Tunisia: Benton et al. 2000) still belong to the same groups recorded in the upper Lower Cretaceous (in Niger and Morocco: Sereno et al. 1994), which implies that the saurischian turnover in Africa, as in Patagonia, took place not earlier than the end of the Cenomanian. In southern Patagonia, Archangelsky (2001) documented a floral turnover in the Aptian, indicated by the disappearance of the Bennettitales, and most representatives of the Cycadales and Ginkgoales. They were replaced by associations dominated by ferns pertaining to the clade Gleicheniaceae. This palaeofloral change, according to Archangelsky (2001), corresponds to the Aptian floral extinction widely recognized (Hallam & Wignall 1997). Archangelsky (2001) suggests that the floral and palaeoenvironmental changes observed in Patagonia may have been due to volcanic activity, rather than to climatic changes produced by sea-level oscillations that have been proposed in other cases. Lamentably, we do not know if a similar floral turnover took place in northern Patagonia at the same time, and we cannot link this event with the

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saurischian turnover recorded in central and northern Patagonia. Probably, the key to understanding this apparent extinction of several saurischian clades, in Patagonia and elsewhere, resides in changes in mid-Cretaceous floras and the disappearance of many plants that were part of the regular diet of sauropods (Salgado 2000). Probably, these floral changes were primarily related to regressive and transgressive cycles, although volcanic activity cannot be discarded (as proposed by Archangelsky 2001 for southern Patagonia). The two major sauropod groups that inhabited Patagonia during the mid-Cretaceous had different masticatory styles, probably correlated with different diets (Calvo 1994). Diplodocoids have slender cylindrical teeth that are restricted to the tip of the snout (Barrett & Upchurch 1994). In contrast, basal titanosauriforms, such as those found in the San Jorge Basin (Martinez 1999), as well as basal titanosaurians recorded in the Neuquen Basin (in the Candeleros Formation and the lowest part of the Huincul Formation), have compressed cone-chisel-like teeth (Calvo 1999, Simon 2001; Simon & Calvo 2002). At least in northern Patagonia, the youngest (upper Cenomanian) undoubted diplodocoid records come from levels lower

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than those containing the oldest remains of titanosaurians with slender teeth that are typical of the uppermost Cretaceous (Fig. 5). The latter forms obviously evolved from groups that were present in the mid-Cretaceous (Apti an-Cenomanian). The occurrence of slender-toothed titanosaurians in the Cenomanian of Africa (Kellner & Mader 1997; Rauhut 1999) suggests that titanosaurians acquired a cylindrical dentition before diplodocoids became extinct, but it is evident that slender-toothed titanosaurians proliferated and diversified after the extinction of the diplodocoids (Salgado 2000). Upchurch (1995) claimed that some groups of sauropods (derived titanosaurians and supposed Upper Cretaceous diplodocoids) survived into the Late Cretaceous by having slender teeth and a distinctive chewing style. However, as already discussed, evidence shows that the main group of sauropods with cylindrical teeth during the Early Cretaceous was the Diplodocoidea (mainly dicraeosaurids and rebbachisaurids), which, according to our interpretation, did not persist into the latest Cretaceous (Fig. 6). In this regard, slender teeth (and a presumably associated chewing style: Barrett & Upchurch 1994) were not intrinsically advantageous, although is unquestionable that only titanosaurians with

Fig. 6. Stratigraphically calibrated cladograms of Cretaceous sauropod and theropod lineages from the Neuquen Basin. A, Sauropods: node 1, diplodocoids; node 2, titanosauriforms. B, Theropods: node 3, carcharodontosaurids; node 4, abelisauroids.

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this particular morphology avoided the mid-Cretaceous extinction. In summary, considering all evidence, it is possible that towards the end of the Early Cretaceous or at the beginning of the Late Cretaceous, there were successive changes in mean temperature (due to unknown factors) and concomitant changes in floral composition. These caused floral changes that subsequently affected the sauropod and theropod faunas. Furthermore, the presumed extinction of the diplodocoids and basal titanosaurians before the Turonian could have created ecological opportunities that led to the dominance of derived titanosaurians (here collectively called eutitanosaurians). The radiation (not the origin) of angiosperms could, in fact, have been linked with the expansion and diversification of titanosaurians with slender peg-like teeth, as Powell (2003) and Salgado (2000) claimed (see Weishampel & Norman 1989 for a similar explanation for the radiation of the ornithopods). Confirmation of this assertion requires more research (see Barrett & Willis 2001 for an extensive discussion about plant-dinosaur coevolution). Certainly, Patagonian mid-Cretaceous angiosperms are scarce, but they were undoubtedly present (Archangelsky 1967; Archangelsky & Gamerro 1967; Romero & Archangelsky 1986), as in other points of the southern Gondwana palynofloral province, although their mere record does not necessarily imply the existence of dinosaurplant coe volution. More problematic are the causes for the extinction of some theropod lineages. But it is reasonable to assume that it was directly due to changes in the sauropod faunas. With the disappearance of carcharodontosaurid theropods, the abelisauroids began to significantly diversify, perhaps stimulated by the radiating eutitanosaurians. In fact, pre-Turonian abelisaurs are scarce in the fossil record, but they become the dominant, if not the exclusive large-sized, theropods during the uppermost Cretaceous (Novas et al. 2005). At present, no diagnostic evidence has been published to support the presence of carcharodontosaurid theropods in deposits younger than Turonian. The association of carcharodontosaurids and abelisauroids during the Early-Late Cretaceous transition, up to Turonian-Coniacian, matches with corresponding changes in sauropod faunas. In the uppermost Cretaceous, the theropod fauna experienced a change comparable to that shown by sauropods, with the disappearance of carcharodontosaurids from the geological record. Some of the supposed carcharodontosaurid remains reported from the higher levels of the

Neuquen Basin (Alcober et al. 1998; Veralli & Calvo 2003) may pertain to large coelurosaurian theropods, similar to those reported from Turonian units (Novas 1998; Coria et al. 2001; Coria & Currie 2002; Porfiri & Calvo 2003) or even in Late Cretaceous levels (Coria & Arcucci 2004). These forms, most of which have yet to be described in detail, retain plesiomorphic features in their skeletons (i.e. dorsal centra with open pleurocoels, large pubic boots) that are present in both carcharodontosaurids and coelurosaurs (Coria & Salgado 1995; Hutchinson & Padian 1997). Coria & Arcucci (2004) report the presence of a large theropod in the Bajo de la Carpa Formation. The size of that specimen is similar to Giganotosaurus, Carcharodontosaurus and other carcharodontosaurids. Nevertheless, the authors caveat that, without derived features, size alone is not enough evidence to justify assignment to Carcharodontosauridae. Consequently, it is possible that the taxonomic diversification shown by titanosaurians (at least in part a response to diplodocoid extinction and, perhaps, to the proliferation of angiosperms) may have also resulted in the replacement of carcharodontosaurids by abelisauroids and coelurosaurs. Even though the available record of dinosaur faunas from the Neuquen Basin is the richest and most diverse succession currently known in the southern hemisphere, there is still much to be learned about the patterns of Patagonian dinosaur evolution. Correlation with other depositional basins in Patagonia in particular and South America overall is essential. Independent chronostratigraphic dating can fine-tune the timing of evolutionary changes, to allow a better understanding of dynamic ecosystems, including both floral and faunal replacements. In this matter, the Neuquen Basin is a unique place to investigate the geobiological evolution of NW Patagonia.

The authors thank the editors for inviting us to participate in this book. PJ. Currie made significant corrections and comments. A. Paulina Carabajal helped in the production of the manuscript. C. Curry-Rogers and M. Lamanna reviewed the manuscript and their comments are greatly appreciated. A. Gerez drew Figure 5. The research was supported by the Museo Carmen Funes Research Funds and the Universidad Nacional del Comahue.

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Index Page numbers in italic denote figure. Page numbers in bold denote tables. Abelisaurus comahuensis 321 Acanthodiscus sp. 60, 64 Acantholissonia gerthi 61 aeolian facies Huitrin Formation 145, 151-152, 157 Troncoso Member 163-164, 167, 168 aeolian systems, flooded 168, 169, 170, 17 174-182 Aeolosaurus 318 Aetostreon 200, 305 Afropollis 76 Agrio Fold and Thrust Belt 3, 16, 18, 29, 30 development 41 stratigraphy 39-40, 40, 42 structure 39, 42-44, 47 uplift Late Cretaceous 43-44 Agrio Formation ammonite biostratigraphy 58, 61, 63, 65, 66, 67 bedding cycles 232, 234-247 calcareous nannofossil biostratigraphy 68, 71, 72 highstand systems tract 154 lithofacies 295, 296, 297, 298-302 marine facies 142-143, 144, 153 organic facies 251-263 palaeoecology 310, 311, 312 palaeoenvironment 309-310, 311, 312-313 palynomorph biostratigraphy 74, 75, 76 stratigraphy, Valanginian—mid-Hauterivian 252-254 Agua de la Mula Member 253, 254 calcareous nannofossils 71 geochemistry 256-257, 256 lateral variation 260-261 lithofacies 297, 298, 312 organic maturation 258 Amargasaurus cazaui 317, 318 ammonite inquilinism 198-199 ammonite zones, early Valanginian-early Hauterivian 60, 69, 70 ammonites biostratigraphy 57-68 correlation with Mediterranean succession 68, 69,70 Los Catutos Member 212, 273, 215, 222 Rio Salado 196, 797 Amphidonte 306, 307, 310, 312 Andean Cordillera 3, 16, 18 stratigraphy 44-45

structure 45-50 Andean Fold and Thrust Belt 37-53 tectonic evolution 50-53 tectonic framework 39 Andes, Neuquen 2, 3, 5, 6 morphostructural units 38 stratigraphy 40 tectonic evolution, 15-32, 37-39, 51 interaction with Neuquen Basin 29-30 Andes, topography 37 Andesaurus delgadoi 318, 320 andesite 21, 23, 26, 42, 44 anoxia see dysoxia-anoxia Aphrodina 199 Aphrodina quintucoensis 302 Aptea notialis 75 Araucariacites australis 74, 75, 76 Araucarioxylon 95, 273-276 arc morphostructural units 38 Arenicolites 193, 196 Argentiniceras noduliferum 62 biozone 58, 61 Asteriacites9Q,9l, 270 Asterosoma 86 92 Auca Mahuida volcano 25, 30 Aucasaurus garridoi 321 Auquilco evaporites 42 Avile Member 141, 253, 298 ammonites 66 calcareous nannofossils 71 Bajocian, fossil coniferous wood 270-276 Balmeiopsis limbatus 76 Barremian, chronostratigraphy 141 basalt alkaline 17, 23, 25, 29 tholeiitic 20 basin back-arc 42 foreland 6, 7, 29 intra-arc 20, 23, 25, 41, 44 batholith Late Cretaceous 29 Patagonian 45 bauplane 285, 286 bay-fill facies, Lajas Formation 87, 88, 90,99 Bayo dome complex 50 bedding cycles shale - marl - limestone

330

INDEX

Cretaceous 231-232, 235-247 geochemistry 235-236, 237, 238 Benioff zone, shallowing 79, 27, 29, 31 benthic associations 198 Berriasella 61 Berriasian-Aptian, biostratigraphy 57-77 'Besaireiceras' australe 64 biopelmicrite 212-216 bioturbation 85-88, 90-94, 93, 98 bivalve association, infaunal 199-200 bivalves Agrio Formation 302, 303, 304, 307 Los Catutos Member 212, 214-215 Vaca Muerta-Chachao Formations 797, 198 body forms see bauplane Bouguer anomalies 47 braidplains, fluvial 112, 114, 118, 779, 14 browsers, mobile 303, 306 bryozoans 307, 308, 372 Burgess sequence 101, 102, 271 Caldera del Agrio 30, 48, 50 Callaqui volcano 50 Callialasporites sp. 74, 76 Callialasporites trilobatus 75 Callianassa 304-305 Campana Mahuida igneous complex 21 Candeleros Formation 28 carbonate, productivity cycles 231-232, 240-241 carbonate facies Agua de la Mula Member 301 Huitrin Formation 42, 144, 148, 149 Mulichinco Formation 117, 725, 126 carnivores mobile epibenthic 303, 306-307 nektonic 303, 308-309 Carnotaurus sastrei 318, 320, 321 Cayanta Formation 21, 42, 44 Caypullisaurus bonapartei 283, 284, 285, 287 290 cement, limestone-marl rhythmite 216, 223 Ceratostreon 200, 306, 307, 310, 312 Cernina fluctuata 306 Cerro Mocho anticline 42, 43 Cerro Parva Negra volcano 25 Chacantuceras ornatum 60, 63, 64, 309 Chachaicosaurus cayi 282, 283, 285, 290 Chachao Formation facies analysis 189, 797, 192 macrofossils 196-200, 202 relative oxygenation 185-203 sedimentation 192 trace fossils 793, 194, 196 Challaco Formation 85, 85, 86, 95 channel-fill 97, 92, 93 Chihuidos ridge 30

Choiyoi Group 40, 40, 43, 44 Chondrites 200, 203 association 793, 194 Chorreado Member 141, 142, 146, 149 Chos Malal fold and thrust belt 41 organic facies variation 251-263, 252 chronostratigraphy 4 Circulodinium distinctum 75, 76 Classopollis sp. 74, 75, 76 Clepsilithus maculosus 71, 72 bioevent 70, 73 climate change, effect on clastic/carbonate cycles 240, 245, 246 Clypeopygus 304 coast, transgressive facies 115, 122-123 coastal plain facies 115, 123 Coenholectypus 306 Cola de Zorro Formation 25, 30, 40, 45, 47-48, 50 Coladas de Fondo de Valle monogenetic field 50 Collipilli Formation 21, 42 Conglomerados de Tralahue 44 Copahue volcano 25, 50 coquinas 298-299, 300 coral, Agua de la Mula 301, 308, 377, 372 Cordillera del Viento 20, 21, 22, 29, 40 basement uplift 42, 44 Cretaceous, Late, uplift, Agrio Fold and Thrust Belt 43-44 Cretaceous, Late-Cenozoic foreland basin phase 3, 6, 1 Cretaceous, Late—Palaeogene arc 21—22, 22 tectonism 29 Cretaceous, Lower aeolian systems 163-182 biostratigraphy 57-77 lithofacies, Agrio Formation 295-302 lowstand 139-160 Mulichinco Formation 109-135 palaeoclimate 245 palaeoecology, Agrio Formation 302-313 sediment supply, Agrio Formation 244-245 Cribroperidinium orthoceras 75, 76 Crioceratites andinus 66, 309 Crioceratites apricus 66 Crioceratites diamantensis 66, 67, 309 biozone 66, 69, 297 Crioceratites schlagintweiti 66, 67 biozone 66, 69, 297 crocodilians 280-281, 282, 283, 285 cross-bedding Lajas Formation 88, 90, 97, 92, 93 Mulichinco Formation 118, 779 Troncoso Member 769 cross-lamination 121, 725, 299-300 Cruciellipsis cuvillieri 71, 72

INDEX

bioevent 70, 71 crust discontinuity 51-52 thin, Loncopue Graben 19-20 Cruziana 120, 121, 123, 125, 299 Cryptoclidus 282, 283, 285, 288, 290 Cucullaea gabrielis 302, 310, 311 Cura Mallm Basin 23, 24, 24, 29, 30, 45 structure 45-47, 50, 51 Cura Mallm Formation 23, 40, 44-45 Cura Niyeu Formation 84, 85, 86, 88, 98, 270 Cura Niyeu-Lajas sequence set 101, 702, 103 Curaco, reworked facies 170, 175, 176, 111, 178-181 Cuyaniceras transgrediens 61 Cuyo Group 43, 84, 85 cycles, orbital 240 cyclicity 231-247, see also rhythmite Cyclusphaera psilata 75, 76 Cymatoceras perstriatum 308, 311 dacite21, 23, 26,42,44 Dactyloidites 90, 93, 270 Dakosaurus sp. 284, 285, 287, 290 Damas-Chaquilvfn structure 50 deformation 26, 28, 29, 30 Agrio Fold and Thrust Belt 30, 41, 42-44 Andean Fold and Thrust Belt 50-51 soft-sediment 179 facies, Troncoso Member 168, 769, 170 delta front facies Lajas Formation 86-88, 87, 89, 99 Mulichinco Formation 114, 118-120, 779 depocentres Cenozoic 45 Late Cretaceous 7 magmatic 23, 45, 46 deposit feeders, burrowing 303, 304-305 deposit-feeders association 200 deposition 7, 8 see also sedimentology deposits, synorogenic 27, 29, 30, 42, 44, 51 diagenesis, limestone 223-224 Diamante Formation 29 Dichadogonyaulax cumula curtospina 75 dilution cycles, clastic 232, 243, 244-245 dinosaurs, mid-Cretaceous 317-324 discontinuity, crustal 51-52 Disparilia sp. 302, 310 dissolution cycles, carbonate 232, 240, 243 drapes, mud 90-91, 97, 92, 779, 120, 121, 722 725 dunes Troncoso Inferior Member 151-152 flooded 163, 167-168, 170, 777 174-182

331

see also aeolian facies dysoxia-anoxia Valanginian - Hauteri vian Agrio Formation 254, 259-261, 298 global 261-263

echinoids Agrio Formation 304 Los Catutos Member 212, 213-214, 275 Eiffelithus primus 11, 72 bioevent 70, 73 Eiffelithus striatus 71, 72 bioevent 70, 73 Eiffelithus windii 71, 72 bioevent 70, 73 embayment facies 90, 116, 123 Entolium 189, 797, 199, 202 Eocene, uplift 29, 50, 51 erg 152 Eriphyla 199 Eriphyla argentina Burckhardt 302 Eryma 307 estuaries 83, 96, 97 facies 115, 121-123, 722 evaporites Agrio Fold and Thrust Belt 42 Huitrin Formation 140, 145, 152-153, 165 exogyrid association 200 extension 26, 29, 30, 40, 45, 46, 47-50 extinction, dinosaur, mid-Cretaceous 322-324 facies associations Agrio-Huitrin Formation 142-153 Troncoso Inferior Member 165-172 Lajas Formation 85-97 Mulichinco Formation 112-126 falling stage systems tract 154, 755, 156, 157 faults Liquine Ofqui 18, 29, 30, 50 fish Chachao Formation 196, 797 pycnodontiform 309, 311 flooding, dunes 170, 174-175, 179, 180, 181 flooding surfaces 97-99, 100, 133 floodplain facies, Lajas Formation 87, 94-96 fluvial facies Huitrin Formation 144, 145, 149-151, 750 Troncoso Member 165, 167 Lajas Formation 87, 94-96, 96, 270 Mulichinco Formation 115, 121 fold and thrust belts 7, 39 see also Agrio Fold and Thrust Belt; Andean Fold and Thrust Belt foraminifera, Los Catutos Member 212, 213, 214, 215 forcing, orbital 240, 245 foreland, migration 21, 27-28 foreland basin phase 6,1

332

INDEX

fossils, trace see ichnofauna Frenguelliceras 61 Gastrochaenolites 88, 795, 196, 301, 312 gastropods Agrio Formation 302 Los Catutos Member 212, 215 Vaca Muerta-Chachao Formations 797, 198 Geosaurus araucanensis 283, 284, 285, 287, 290 Gervillaria alatior 304, 307, 311 Gervillella aviculoides 305 Gigantosaurus carolinii 321 Glossifungites 97 Gondwana Orogen, collapse 5 Gondwana, western margin biostratigraphy 5 evolution 759 tectonics 2-3, 5 Gordia 793, 196 granitoids 20-21, 29 Groebericeras 59, 61 growth-ring analysis 277, 274-275 guild analysis, Agrio Formation 302, 303, 304-309 Haqius circumradiatus 71 Hauterivian, index ammonites 65, 67 hemicycles carbonate 236 clastic 235 highstand systems tracts 97, 132 Holcoptychites agrioensis 60, 64, 65, 309, 377 Holcoptychites compressum 60, 64 Holcoptychites magdalenae 64 Holcoptychites neuquensis 64, 65 biozone 60, 64, 69, 296, 304 Hoplitocrioceras gentilii 66, 67, 309 biozone 60, 66, 69, 296 Hoplitocrioceras giovinei 66, 67 Hoploparia 307 Huincul Fault Zone 29 Huitrin Formation 40, 42, 139-140, 747, 143, 146, 147 facies associations 142-153, 144, 145, 746 lowstand wedges 139-140, 153-160 master sequence boundary 157-158 sequence stratigraphy 154-158 transgressive systems tract 158 see also Troncoso Formation hummocks 299-300 hydrocarbons 5, 217, 253, 254, 257-259 Agrio Formation 251, 253 Hystrichosphaerina neuquina 74, 75 ichnofauna Agrio Formation 298, 299

Lajas Formation 86, 88, 90-92, 95, 98, 702 Mulichinco Formation 121 as proxy for relative oxygenation 185 Rio Salado 192, 793, 194, 795, 196, 200-202 ichthyosaurs 280-281, 282, 283, 285 llokelesia aguadagrandensis 320 Inoceramus 236, 239, 305 inversion, tectonic 110 Isognomon lotenoensis 307 Isognomon ricordeanus 304, 307, 377 isotopes black shales, 513C anomaly 263 Los Catutos rhythmite 216, 278, 279 222-224 Jurassic Early-Early Cretaceous arc volcanism 20-21 biostratigraphy 7-8 post-rift phase 2-3, 5, 6, 7 sequence stratigraphy 8-9 tectonism29, 50-51 evaporites, Agrio Fold and Thrust Belt 42 Late-Early Cretaceous, palaeo-oxygenation studies 185-203 marine reptiles 279-291 Middle Lajas Formation fossil coniferous wood 267-276 sedimentology 83-104 stratigraphy 85 kaolinite 216, 220 Karakaschiceras attenuatus 60,61, 63, 132, 133, 306 kerogen 217, 220, 257-258 Komplott sequence 102, 10 La Tosca Member 141 lagoon facies 90, 116, 123 Lajas Formation 85, 86 facies associations 85-97 facies model 96 fossil coniferous wood 267-276 geological setting 270-271 taxonomy 273-274 sedimentology 85-96, 87 sequence stratigraphy 96-104, 269 boundaries 101 deltaic successions 99 flooding surfaces 97-99 heterolithic tidal successions 99 model 103-104

INDEX

parasequences 97, 99, 100, 269 tidal channel successions 99, 270 Laplatasaurus 318 Las Damas valley 47, 48 Leptosalenia 306, 312 Ligabueino andesi 320 Limaysaurus 319 limestone lithographic 207 micritic see shale-marl-limestone rhythmic bedding limestone-marl rhythmite, Los Catutos 207-226 Liopleurodon sp. 283, 284, 285, 287, 290 Liquifie Ofqui Fault Zone 18, 29, 30, 50 Lissonia riveroi 61, 62 biozone 60, 61, 69, 112, 132 Lithophaga 304 lithostratigraphy 4 Lithraphidites bollii 11,72 bioevent 70, 73 Loncopue Graben 3, 15-16, 17, 18, 23, 25, 30 crustal thinning 19-20 Los Catutos Member limestone-marl rhythmite 207, 210-226 burial history 209, 224, 225, 226 chemistry 216 clay minerals 216, 220-221 diagenesis 223-224 isotopic composition 216, 218, 219 222-224 organic matter 217 origin 217, 219 palaeoenvironment 221-222 petrography 212-216 Los Molles Formation 84, 85, 86 lowstand, Lajas formation 103 lowstand systems tract 109—110 Huitrin Formation 157-158 Mulichinco Formation 129-132 lowstand wedge Huitrin Formation 139-140, 141, 158-160 evolution 153-154 sequence stratigraphy 154-158 Mulichinco Formation 109-110, 777, 130 133-134 Lyticoceras pseudoregale 64 macrofossils, Vaca Muerta-Chachao Formations 196-200, 207, 202 Macromesodon agrioensis 309 magmatism, arc 15, 16, 20, 46 Cretaceous, Late-Palaeogene 21-22, 22 Jurassic-Early Cretaceous 20-21 migration 22, 27, 29 Miocene, middle-late 23-24 Oligocene-early Miocene 22-23, 26 Pliocene-Pleistocene 25

333

styles 25-28 Maresaurus coccai 280, 282, 285, 288, 290 marine facies, Agrio Formation 142-143, 144, 148 marl see limestone-marl rhythmite; shalemarl-limestone rhythmic bedding Mendoza Group 40, 42, 57, 58 Mendoza Shelf 186 facies analysis 189-192 sedimentation 192 stratigraphy 787, 188 Mesozoic, Neuquen Basin deposits 39-40, 42 Metriorhynchus casamiquelai 288 Metriorhynchus potens 290 Meyerella rapax 306 Micrantholithus hoschulzii 71, 72 micrite 241 microcarnivores, cemented 303, 308 Milankovitch cycles 240, 245 Mimachlamys robinaldina 307, 377 Miocene, middle-late arc 23-24 tectonism 30, 51 Mitrauquen Formation 23, 45 Modiolus 85 Modiolus cf. subsimplex 305 Mollesaums perihallus 282, 283, 285, 288 mouth bars 96 Muderongia brachialis 76 Muderongia staurota 75, 76 mudflats 2, 9, 93 mudstone Agrio Formation 299-300 Lajas Formation 85-86, 88, 92, 95 Mulichinco Formation 121, 123, 126 Mulichinco Formation ammonite biostratigraphy 58, 61 calcareous nannofossil biostratigraphy 68, 71 facies associations 112, 114-117, 118-26, foldout geology 110-111, 773 highstand systems tract 132-133 lowstand systems tract 129-132 lowstand wedge 109-110, 777, 130 133-134 palaeoflow 120, 123, 124 palaeogeographic evolution 128, 131-132 palynomorph biostratigraphy 74 sequence stratigraphy 126, 727, 129 transgressive systems tract 132 Muraenosaurus sp. 282, 283, 285, 288, 290 Myoconcha transatlantica 304, 305, 377 Mytilus 307 nannoconids 74 Nannoconus bucheri 72, 74 Nannoconus circularis 74 Nannoconus ligius 11,72

334

INDEX

bioevent 70, 73-74 nannofossils, calcareous bioevents 70, 71, 73-74 biostratigraphy 68-71, 72 nekton 196-197,202 Neocomites sp. 60, 61, 63, 64 Neocomites wichmanni 62 biozone58, 60, 61,69, 772 Neocosmoceras sp. 61, 71 Neogene-Quaternary, Andean Cordillera 47-50 Neohoploceras 61 Neuquen Basin evolution 2-3, 5, 6, 7, 164-165 geological setting 1, 2, 3, 84-85, 140-142, 209-210, 232-234 palaeogeography 186-188 stratigraphy 4, 40 Neuquen Embayment 2, 3, 23, 25, 38, 186, 787 234 sedimentation 192 Neuquensaurus 318 Neusticemys neuquina 283, 284, 285, 287, 290 Nevados de Chilian Volcanic Group 23 North Patagonian Massif 2, 244-245, 244 Notoemys laticentralis 283, 284, 285, 287, 290 Nucleolites 304 offshore shelf facies 85-86, 87 Olcostephanus (Jeannoticeras) agrioensis 64 Olcostephanus (Olcostephanus) atherstoni 61, 63, 309 biozone 60, 61, 69, 772, 132 Olcostephanus (Olcostephanus) boesei 64 Olcostephanus (Olcostephanus) laticosta 60, 64, 65, 309 Olcostephanus (Olcostephanus) leanzai 60, 64 Olcostephanus (Olcostephanus) variegatus 66 Olcostephanus (Viluceras) permolestus 60, 61, 63, 133 Oligocene-early Miocene arc 22-23, 24, 26, 44-45 tectonism 29-30, 51 Oligosphaeridium complex 75, 76 Oosterella 64 Ophiomorpha 86, 92 Ophthalmosaurus sp. 282, 283, 285, 290 organic facies Agrio formation 251-263 geochemistry 256-257 ostracods, Los Catutos Member 212, 213 Owl sequence 702, 103, 271 oxygenation relative 185-203 from ichnofauna 200-202, 207 from body fauna 207, 202 oysters, cemented 117, 725, 126, 192, 797

palaeobiology 9-10, 267-276, 279-291, 295-313, 317-324 palaeoclimate, Lower Cretaceous, Agrio Formation 245, 246 palaeoenvironment, Jurassic marine reptiles 285, 287-288 Palaeogene, Cura Mallin Basin 45-47 Palaeophycus 86, 92, 121, 143, 793, 196, 236 palynomorph biostratigraphy 74, 75, 76 Pampa de Tril calcareous nannofossils 71 reworked facies 170, 777, 172, 773 174-175, 178-181 Panopea gurgitis 302 Parahaentzschelinia 87 parasequences 97, 99, 700, 131, 132 Paraspiticeras groeberi 67 biozone 66, 68, 69, 297, 372 Parsimonia antiquata 307-308, 370 Patagonian Batholith 45 pelbiomicrite 212-216, 214 Pellegrinisaurus 318 pellets 212, 216 Pholadomya gigantea 302 Phycodes 193, 196 Pichaihue syncline 43, 44 Pichi Tril Andesite 23 Pilmatue Member 253-263 ammonites 61, 66 calcareous nannofossils 71 geochemistry 256-257, 256 lateral variation 258-260 lithofacies 296, 298, 370, 377 organic maturation 258 Pinna robinaldina 305, 377 planktonic associations 797, 198 Planolites 90, 92, 93, 123, 793, 196, 236 Plattenkalke 207 plesiosaurs 280, 283, 285, 309 Pleurotomaria gerthi 306 Plicatula 307 Pliocene-Pleistocene arc 25 tectonism 30, 51 pliosaurs 280, 282, 283, 285, 287 post-rift phase 5, 6,1 Principal Cordillera 17, 23, 25 granitoids 20-21 prodelta facies 85-86, 87, 96, 99, 115, 120-121 productivity cycles, carbonate 232, 240-241 progradation, Pilmatue Member 253 Protaxius 304, 305 proto-Pacific ocean 5, 6, 9 Protocallianassa 304, 305 Protohemichenopus neuquensis 261, 302, 370 Pseudofavrella angulatiformis 63, 64, 370 biozone 60, 64, 69, 772, 133, 296, 370 Pseudofavrella garatei 64

INDEX

pseudoplanktonic associations 797, 198 pterosaurs 283 Pterospermella australiensis 75, 76 Pterotrigonia coihuicoensis 311 Ptychomya koeneni Behrendsen 302, 311 Puesto Burgos ignimbrites 44 Purranisaurus potens 283, 284 Pushme-Fullyou sequence 101, 102 Pycnodontiformes 309, 311 Pygorhynchus 304 radiolarians, Los Catutos Member 212, 214 ramp, carbonate 149, 192 see also carbonate facies Rayoso Formation 28, 58 Rayoso group 40, 57, 58 reptiles, marine Jurassic 279-291 bauplane 285, 286 palaeoenvironment 285, 287-288 palaeogeographic distribution 288, 290 retro-arc morphostructural units 26, 38, 40, 41, 44,46 retro-arc system 5, 7 reworked facies, Troncoso Member 769, 170, 172 rhaxes, Los Catutos Member 212, 222 Rhizocorallium 193, 196 rhythmite, limestone-marl Tithonian 207-226 burial history 224, 225, 226 origin 217, 219 palaeoenvironment 221-222 see also bedding cycles rifting 6, 29, 51-52 Rio Picunleo 47, 48-50 Rio Salado see Salado River valley ripples 121, 123, 124, 299-300 flooded aeolian systems 769, 170, 172 Rocasaurus 318 roll-back velocity 21, 29, 38 Rosselia 86, 92 Rotularia 199, 202, 203 Salado River valley 787, 188, 189, 790 macrofossils 196-200 trace fossils 192, 793, 194, 795, 196 Saltasaurus 318 sandflats 93-94 sandstone Agrio Formation 298-300 Huitrin Formation 143, 148, 151 Lajas Formation 87-88, 90, 91, 92 Mulichinco Formation 118, 119, 120, 121 see also Avile Member Sarcinella occidentalis 308 sauropoda 317-318 scavengers 303, 308

335

Schaubcylindrichnus 86 Scolicia 90, 91, 270 sea-level change 8, 9, 97 Barremian-Aptian, Huitrin Formation 139-140, 156, 158, 759, 160 Valanginian Agrio Formation 42, 253 Mulichinco Formation 110, 126, 737 sediment, cyclicity 231-2 sediment supply, Lower Cretaceous, Agrio Formation 244-245, 246 sedimentology Jurassic, Lajas Formation 83-104 Tithonian-Valanginian, Mendoza Shelf 188-192 sequence boundaries 101 Serie Andesitica 42, 44 serpulids 797, 199, 307-308 shale, black Agrio Formation 111, 141, 251, 253, 254, 256, 256-61, 263, 298, see also Spitidiscus shale Vaca Muerta Formation 189, 202-203 shale, grey Agrio Formation 298-299 Chachao Formation 189 shale-marl-limestone rhythmic bedding, Agrio Formation 235-247 shelf facies 116, 124-126, 725 shoreface facies 116, 124, 725 Huitrin Formation 143, 144, 147-149, 148 Sierra de Trocoman 47, 48 Sierra del Chacaico 85, 86, 96, 267 Sierra Pintada Massif 2 siltstone, Lajas Formation 86, 87 Skolithos 123, 124, 125 soft-sediment see deformation, soft-sediment Speetonia colligata 71, 72 Sphaera koeneni 302 Spiticeras damesi 61, 62, 74 biozone 58, 61 Spitidiscus riccardii 67, 298, 309 biozone 66, 69, 297 Spitidiscus shale 66, 253, 254, 255, 256, 298 geochemistry 256-257 organic maturation 258 sponges Agrio Formation 308, 312 Los Catutos Member 212 Steinmanella 199 Steinmanella pehuenmapuensis Leanza 302, 310 Steinmanella transitoria (Steinman) 302, 377 Steneosaurus gerthi 280 Stenopterygius grandis 282, 283 stratigraphy, sequence 8-9 Lajas Formation 96-104 Mulichinco Formation 126, 727, 129 subduction 3, 5, 6, 26

336

INDEX

turtles 283, 285 angle 7, 27, 28, 30, 31 Tylostoma jaworskii 306 Substeueroceras koeneni zone 60-61 suspension feeders 199-200, 302, 303, 304, 305, 307-308 swales 300 uplift synrift phase 5, 6 Eocene 50, 51 Late Cretaceous 29, 43-44 up welling, asthenospheric 25-26, 28, 29 tectonics history 2-3, 4 interaction of Andes and Neuquen Basin Vaca Muerta Formation 29-30, 31, 37-39 ammonite biostratigraphy 57-58, 60, 61, 62 styles 25-28 calcareous nannofossil biostratigraphy 68, 71, Teichichnus 92, 121, 123, 193, 196, 236 72 Teredolites 90, 270 facies analysis 189, 191,192 Thalassinoides 86, 92, 93, 194-196, macrofossils 196-200 200-202, 203, 236, 239, 254, 305 marine reptiles 283, 287 association 193, 194 palynomorph biostratigraphy 74-76 theropoda 318, 320-322 relative oxygenation 185-203 'Thurmanniceras' 61 sedimentation 192 'Thurmannites pertransiens Sayn' 61 trace fossils 193, 194, 196 tidal channel facies Valanginian Lajas Formation 87, 90-91, 91, 92, 96, 98, 99, early-early Hauterivian, ammonite zones 60 270 index ammonites 62, 63 Mulichinco Formation 114, 120 Mulichinco Formation 110-111 tidal flats facies, Lajas Formation 87, 92-94, 96, biostratigraphy 772 99 Valanginian - mid-Hauteri vian titanosaurs 318 Agrio Formation 251-263 Tithonian, Middle, rhythmite 207-226 transgression 253 Tordillo Formation 40, 43 Valanginites argentinicus 61 Tralalhue Conglomerates 30 volcanism, arc transgression 7, 8, 179, 186, 253 Cretaceous, Late-Palaeogene 21-22 transgressive systems tracts 97, 101, 104, 109, Jurassic-Early Cretaceous 20-21 132, 158 Miocene, middle-late 23-24 transpression 22, 29 Oligocene-early Miocene 22-23, 26, Trapa-Trapa Formation 23 144-45 Tres Chorros extensional system 42, 47 Pliocene-Pleistocene 17, 18, 25, 45 Triassic, Late-Early Jurassic synrift phase 2-3,5,6 Trigonia carinata Agassiz 302, 311 Watznaueria 71 Trolon volcano 50 Weavericeras vacaensis 67 Trolope dome complex 50 biozone 60, 66, 69, 296 Trolope valley 50 wood, fossil 95 Tromen volcano 30 coniferous Troncoso Member 141-142,146,147, 150, 151, Lajas Formation 267-276 152 geological setting 270-271 depositional model 177-182 taxonomy 273-274 facies associations 165-172 sedimentary architecture 172-177 soft-sediment deformation 163-182 Trypanites 193, 196 Xenotarsosaurus bonapartei 320