Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and diagenetic reference models (Geological Society Special Publication 329)

Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models

The Geological Society of London Books Editorial Committee Chief Editor

BOB PANKHURST (UK) Society Books Editors

JOHN GREGORY (UK) JIM GRIFFITHS (UK) JOHN HOWE (UK) RICK LAW (USA) PHIL LEAT (UK) NICK ROBINS (UK) RANDELL STEPHENSON (UK) Society Books Advisors

MIKE BROWN (USA) ERIC BUFFETAUT (FRANCE ) JONATHAN CRAIG (ITALY ) RETO GIERE´ (GERMANY ) TOM MC CANN (GERMANY ) DOUG STEAD (CANADA ) MAARTEN DE WIT (SOUTH AFRICA )

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 Book Editors 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 book for the Society can be found on its 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: VAN BUCHEM , F. S. P, GERDES , K. D. & ESTEBAN , M. (eds) 2010. Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329.

GRE´ LAUD , C., RAZIN , P. & HOMEWOOD , P. 2010. Channelized systems in an inner carbonate platform setting: differentiation between incisions and tidal channels (Natih Formation, Late Cretaceous, Oman). In: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 163–186.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 329

Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models

EDITED BY

F. S. P. VAN BUCHEM Maersk Oil Qatar AS, Qatar

K. D. GERDES Shell International Exploration & Production, The Netherlands

and M. ESTEBAN Repsol/YPF, Spain

2010 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 over 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 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. þ44 (0)20 7434 9944; Fax þ44 (0)20 7439 8975; E-mail: [email protected]. 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. þ44 (0)1225 445046, Fax þ44 (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 2010. 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 Ltd, Saffron House, 6– 10 Kirby Street, London EC1N 8TS, UK. Users registered with the Copyright Clearance Center, 222 Rosewood Drive, Danvers, MA 01923, USA: the item-fee code for this publication is 0305-8719/10/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN 978-1-86239-292-2 Typeset by Techset Composition Ltd, Salisbury, UK Printed by MPG Books Ltd, Bodmin, UK Distributors North America For trade and institutional orders: The Geological Society, c/o AIDC, 82 Winter Sport Lane, Williston, VT 05495, USA Orders: Tel. þ1 800-972-9892 Fax þ1 802-864-7626 E-mail: [email protected] For individual and corporate orders: AAPG Bookstore, PO Box 979, Tulsa, OK 74101-0979, USA Orders: Tel. þ1 918-584-2555 Fax þ1 918-560-2652 E-mail: [email protected] Website: http://bookstore.aapg.org India Affiliated East-West Press Private Ltd, Marketing Division, G-1/16 Ansari Road, Darya Ganj, New Delhi 110 002, India Orders: Tel. þ91 11 2327-9113/2326-4180 Fax þ91 11 2326-0538 E-mail: [email protected]

Contents VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. Mesozoic and Cenozoic carbonate systems of the Mediterranean and the Middle East: stratigraphic and diagenetic reference models – an introduction

1

GERDES , K. D.,WINEFIELD , P., SIMMONS , M. D. & VAN OOSTERHOUT , C. The influence of basin architecture and eustacy on the evolution of Tethyan Mesozoic and Cenozoic carbonate sequences

9

LUCˇ IC´ , D., IVKOVIC˙ , Z., FORSˇ EK , G., TAKACˇ , D., BUBNIC˙ , J. & KOCH , G. Depositional sequences and palynology of Triassic carbonate-evaporite platform deposits in the Palmyrides, Syria

43

PIERRE , A., DURLET , C., RAZIN , P. & CHELLAI , E. H. Spatial and temporal distribution of ooids along a Jurassic carbonate ramp: Amellago outcrop transect, High-Atlas, Morocco

65

AURELL , M., BA´ DENAS , B., IPAS , J. & RAMAJO , J. Sedimentary evolution of an Upper Jurassic epeiric carbonate ramp, Iberian Basin, NE Spain

89

EMBRY , J.-C., VENNIN , E., VAN BUCHEM , F. S. P., SCHROEDER , R., PIERRE , C. & AURELL , M. Sequence stratigraphy and carbon isotope stratigraphy of an Aptian mixed carbonate-siliciclastic platform to basin transition (Galve sub-basin, NE Spain)

113

DROSTE , H. High-resolution seismic stratigraphy of the Shu’aiba and Natih formations in the Sultanate of Oman: implications for Cretaceous epeiric carbonate platform systems

145

GRE´ LAUD , C., RAZIN , P. & HOMEWOOD , P. Channelized systems in an inner carbonate platform setting: differentiation between incisions and tidal channels (Natih Formation, Late Cretaceous, Oman)

163

RAZIN , P., TAATI , F. & VAN BUCHEM , F. S. P. Sequence stratigraphy of Cenomanian –Turonian carbonate platform margins (Sarvak Formation) in the High Zagros, SW Iran: an outcrop reference model for the Arabian Plate

187

VAN BUCHEM , F. S. P., ALLAN , T. L., LAURSEN , G. V., LOTFPOUR , M., MOALLEMI , A., MONIBI , S., MOTIEI , H., PICKARD , N. A. H., TAHMASBI , A. R., VEDRENNE , V. & VINCENT , B. Regional stratigraphic architecture and reservoir types of the Oligo-Miocene deposits in the Dezful Embayment (Asmari and Pabdeh Formations) SW Iran

219

JANSON , X., VAN BUCHEM , F. S. P., DROMART , G., EICHENSEER , H. T., DELLAMONICA , X., BOICHARD , R., BONNAFFE , F. & EBERLI , G. Architecture and facies differentiation within a Middle Miocene carbonate platform, Ermenek, Mut Basin, southern Turkey

265

WEIDLICH , O. Meteoric diagenesis in carbonates below karst unconformities: heterogeneity and control factors

291

ROSALES , I. & PE´ REZ -GARCI´ A , A. Porosity development, diagenesis and basin modelling of a Lower Cretaceous (Albian) carbonate platform from northern Spain

317

vi

CONTENTS

SHARP , I., GILLESPIE , P., MORSALNEZHAD , D., TABERNER , C., KARPUZ , R., VERGE´ S , J., HORBURY , A., PICKARD , N., GARLAND , J. & HUNT , D. Stratigraphic architecture and fracture-controlled dolomitization of the Cretaceous Khami and Bangestan groups: an outcrop case study, Zagros Mountains, Iran

343

RONCHI , P., DI G IULIO , A., CERIANI , A. & SCOTTI , P. Contrasting fluid events giving rise to apparently similar diagenetic products; late-stage dolomite cements from the Southern Alps and central Apennines, Italy

397

Index

415

Mesozoic and Cenozoic carbonate systems of the Mediterranean and the Middle East: stratigraphic and diagenetic reference models – an introduction F. S. P. VAN BUCHEM1*, K. D. GERDES2 & M. ESTEBAN3 1

Maersk Oil Qatar AS, P.O. Box 22050, Doha, Qatar

2

Shell International Exploration and Production bv, Kessler Park l, 2288 GS Rijswijk, The Netherlands 3

Repsol/YPF, Paseo de la Castellana 280, Madrid, 28046, Spain

*Corresponding author (e-mail: [email protected])

The contributions in this volume originally formed a set of presentations at a conference on the same theme held in Mallorca, Spain in 2006. The goal of this conference was to investigate the potential to develop age or architecture specific reference models for carbonate systems and reservoirs similar to those successfully developed for siliciclastic systems. The conference focused on the Mesozoic and Cenozoic carbonate sequences of the Mediterranean and Middle East. These sequences were chosen for a number of reasons. Firstly, they represent sequence development in a variety of basin settings within a contiguous geographical entity, the former NeoTethys Ocean (Fig. 1). The sequences were also formed predominantly within tropical or sub-tropical climatic zones (cf. Schlager 2003). Finally, the high levels of industry and academic interest in the region have generated many excellent multidisciplinary studies of these sequences, based on both the comprehensive datasets of hydrocarbon-bearing strata and the excellent surface exposures in the region. In general, all Earth models underestimate the complexity of the subsurface and hence are intrinsically inaccurate. The value of developing such models, however, lies in the improved understanding of the processes controlling sequence development gained from their application (e.g. Ahr 1973; Read 1985; Burchette & Wright 1992; Handford & Loucks 1993; Pomar 2001; Bosence 2005). Extrapolating from data rich examples into areas where data coverage is poorer obliges us to distil out the generic from the specific and to propose appropriate subsurface analogues. The two key variables in the development of carbonate sequences and reservoirs, which have a negligible effect on siliciclastic system development, are the biological origin and unstable chemical nature of the constituent material. Carbonate sequences are generated by dynamic, living systems that are biologically reactive and which frequently

modify in response to changes in depositional environment. These responses can be both short term, for example when organisms react to locally induced palaeoecological changes in a particular depositional setting, or long term, as carbonate producing organisms evolve in response to physical changes in the Earth’s biosphere (e.g. James 1983; Schlager 1991; James & Bourque 1992; Ager 1993; Kiessling et al. 1999, 2002; Simmons et al. 2007; Markello et al. 2008; Pomar & Hallock 2008; Pomar & Kendall 2008). A challenge for Earth scientists is to assess how the interaction of global processes (sea level, climate, plate tectonics, global carbon budget) along with biological evolution, determined the nature of both local and regional carbonate factories; and whether this interaction created an infinite variety of carbonate depositional systems, or a more limited and predictable number of sedimentation patterns distributed in a systematic manner in time and space. Similarly, the chemical reactivity of carbonate depositional systems introduces a complexity that is largely absent in quartz-dominated siliciclastic systems. The chemically reactive nature of carbonates is evident in the dissolution and precipitation of different mineral species in both the early phases of deposition (e.g. Choquette & James 1983; Budd et al. 1995) and in later burial and/or inversion phases controlled by geodynamic processes (e.g. Horbury & Robinson 1993; Purser et al. 1994; Braithwaite et al. 2004). This chemical reactivity commonly results in important alterations of the porosity/permeability patterns imposed by primary depositional textures and constituents. Similarly, the challenge here is to establish whether these processes generate an infinite variety of diagenetic end products or if certain generic and predictable patterns can be distinguished. Most contributions in this volume deal with the stratigraphic architecture of carbonate depositional systems, whereas a smaller number of case studies also include the diagenetic aspects.

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 1–7. DOI: 10.1144/SP329.1 0305-8719/10/$15.00 # The Geological Society of London 2010.

2 F. S. P. VAN BUCHEM ET AL. Fig. 1. Geographical map of the study area showing the locations of the case studies presented in this book. 1, Lucic et al.; 2, Pierre et al.; 3, Aurell et al.; 4, Embry et al.; 5, Droste; 6, Grelaud et al.; 7, Razin et al.; 8, van Buchem et al.; 9, Janson et al.; 10, Weidlich; 11, Rosales et al.; 12, Sharp et al.; 13, Ronchi et al.

MESOZOIC AND CENOZOIC CARBONATE SYSTEMS OF THE MEDITERRANEAN

Stratigraphic reference models Carbonate-producing organisms are highly sensitive to changes in depositional environment, which makes them more precise recorders of physical and chemical, global- and regional-scale changes than siliciclastic systems. The biological evolution induced successive replacements of the main carbonate production modes, both in biotically-controlled skeletal production and biotically-induced precipitation (sensu Schlager 2003). Consequently, carbonate sequences show age specific stratigraphic characteristics (e.g. Kiessling et al. 2002; Markello et al. 2008), which suggest that any reference model classification of carbonate sequences should have a chronological aspect. A key property of the hydrocarbon prospectivity of certain carbonate sequences is their capacity to be self-sourcing. This occurs when marine source rocks (organic carbon) are deposited as the basinal equivalents of shallow-water carbonates (mineral carbon) with reservoir potential within the same sequence. These intervals are of obvious economic interest and appear to be confined to specific intervals of the geological record (Tissot 1979; Klemme & Ulmishek 1990; Greenlee & Lehman 1993; van Buchem et al. 2005). Several examples of self-sourcing carbonate hydrocarbon systems are present in the area of interest. These were deposited during the Middle Triassic, Late Jurassic, Middle Cretaceous, Palaeogene and Oligocene – Miocene periods, and the industry interest is reflected in the comprehensive datasets and studies available for these intervals. The introductory contribution by Gerdes et al. illustrates how changes in regional tectonics, eustacy and basin architecture acted as first-order controls on the variation of accommodation space available for carbonate deposition through time. The structural controls evident in the deposition of carbonate sequences are illustrated to explain why strikingly different carbonate platforms are developed during the same stratigraphic time interval in adjacent regions. The paper also illustrates, through a series of maps, the expansion of accommodation space that occurred during the Jurassic period and that culminated in the Late Cretaceous; and the even more rapid destruction of accommodation space during the Cenozoic in the Tethyan realm. The contrasts in carbonate sequence development both spatially and temporally are also related to variations in the surfaces of deposition that were ultimately controlled by the evolving submarine basin architectures. The Triassic carbonate sequences in the study area illustrate both the relatively rapid recovery of carbonate organisms after the mass extinction event at the Permo-Triassic boundary and the control of

3

basin architecture on carbonate sequence development. Two strikingly different systems are developed on differing basin architectures: (a) Very large, shallow water carbonate platforms characteristic of the post-rifted passive margin of the Arabian Plate (e.g. Sharland et al. 2001). These are typically organized in decametrescale depositional sequences consisting of carbonates, evaporites and dolomites. Lucic et al. provide an overview of the Triassic hypersaline ramp sequences of Syria that are an example of this type of platform deposition. The lithological variability is demonstrated in a framework of third-order depositional sequences, constrained by new palynological age dating. (b) Isolated buildups of limited areal size, but of great thickness, that are characteristic of synrift deposition on the rapidly subsiding tilted fault blocks created during active extension along the northern NeoTethys margins. These are now spectacularly exposed in the Dolomites of northern Italy and are extensively documented in the literature (e.g. Bosellini 1984; Maurer 2000; Brack et al. 2007). The Jurassic carbonate sequences are characterized by a variety of depositional geometries and carbonate producers, including microbial mounds, oolitic and bioclastic ramp systems (e.g. Leinfelder et al. 1994; Kiessling et al. 2002). The platform type that has attracted most attention from the petroleum industry in the focus area is the oolitic ramp. Two well-illustrated examples are presented in this volume. Pierre et al. describe a Lower –Middle Jurassic oolitic ramp system in the Atlas Mountains of Morocco, which has continuous exposure over a distance of 37 km, and has been dated precisely with ammonites. This paper describes a suite of fourth-order sequences that can be followed over tens of kilometres. An important feature of this outcrop study is the observation of beds of muddy, low-relief, transgressive deposits alternating with grainy, oolitic, high relief, higher energy regressive deposits within the fourth-order sequences. This clear organization is attributed to an as yet poorly understood climatic –biota interaction. Aurell et al. present a comprehensive study of the Upper Jurassic ramp systems in NE Spain along a 200 km long transect. This analysis documents shallow to deep water lithological transitions of oolitic, bioclastic and microbial-algal facies. The sequence stratigraphy is chronologically constrained by ammonite dating and is summarized in a sequence framework of first to fifth-order sequences. The Cretaceous carbonate sequences in the study area are characterized by large-scale carbonate platforms, illustrating high frequency variations in carbonate margin growth and the contemporaneous

4

F. S. P. VAN BUCHEM ET AL.

deposition of extensive organic-rich shales in the oceans and intrashelf basins. Two time intervals, the Barremian/Aptian and Cenomanian/Turonian, have attracted particular attention due to their hydrocarbon content and as suitable sections with which to study the effect and extent of global events such as oceanic anoxic events (OAEs) on carbonate deposition. Barremian/Aptian times are characterized by the deposition of specific depositional facies that can be traced along the NeoTethyan coastlines from the Mediterranean to the southern margin of the Arabian Plate. These grainy, rudist-dominated facies of the ‘Urgonian’ platforms, contain the characteristic ‘Orbitolina Beds’ and the Lithocodium/Baccinella microbial facies, and pass laterally into the organic-rich basinal deposits of the Aptian OAE la or Selli event (e.g. Menegatti et al. 1998; Pittet et al. 2002; Wissler et al. 2002; Weissert & Erba 2004; Follmi et al. 2006; van Buchem et al. 2009). Embry et al. provide a good example of this facies succession exposed in excellent outcrops that illustrate a full carbonate sequence from platform to basinal setting along a 16 km transect in NE Spain. A particular aspect of this study is the description of well-preserved, Upper Aptian shallow-water deposits. These facies, which are absent in most peri-Tethyan localities, provide an example of an almost continuous Aptian carbonate sequence in outcrop. Droste provides a well documented subsurface study of the Aptian shallow-water platforms of Oman, based on a comprehensive dataset of 3D and 2D seismic data, well logs and core material. The similarity in the overall depositional architecture and facies distribution between these two localities, which are 5000 km apart, is striking, and makes a strong case for proposing an Aptian rimmed platform/intrashelf basin reference model (see also van Buchem et al. 2010). The Cenomanian/Turonian platforms of the Arabian plate are presented in three case studies: both Droste and Grelaud et al. present subsurface and outcrop examples from Oman, and Razin et al. present a study of a fully exposed sequence along a 10 km transect in SW Iran. These papers provide complete and detailed pictures of the stratigraphic architecture that is considered typical of carbonate deposition on the margins of NeoTethys during this time interval. These authors describe a stratigraphic organization of four to five third-order depositional sequences, two of which include organic-rich source horizons in intrashelf basinal settings. The subtle variation in carbonate facies is thought to be in response to increased rates of sea level rise. These sequences are considered to have been initially eustatically driven and illustrate well-developed Mid-Cenomanian incised valley systems. Carbonate deposition became tectonically

influenced during the late Cenomanian and Turonian, as obduction around the margins of the Arabian Plate caused regional inversion that was followed by another phase of surface incision. This tectonostratigraphic pattern is documented by the integration of the mapping of world-class outcrops, the interpretation of 3D seismic data, core descriptions and well log interpretations. These examples provide good sample sets upon which to base a Cenomanian/Turonian high angle rimmed platform/intra-shelf basin reference model for the Arabian Plate. Carbonate sequences deposited during the Palaeogene and Neogene range from large Eocene platforms in a foreland basin setting, dominated by Nummulites, to smaller, more complex Oligocene and Miocene carbonate systems developed on submerged extensional and thrusted fault block crests (e.g. Esteban 1996; Franseen et al. 1996; Pomar 2001; Kelling et al. 2005; Gerdes et al. 2010) This variation in sequence type was influenced by a combination of structural control, depositional substrate morphology and the climatic changes associated with the passage from the Palaeogene thermal maximum, into the ice-house conditions that commenced at the Eocene–Oligocene boundary (Miller et al. 2005). The high amplitude, glacio-eustatically driven sea-level fluctuations typical for this time interval, had significant implications for the depositional geometries developed and led to the temporary isolation (and periodic desiccation) of the intramontane/intraplate basins in the area of interest such as the Dezful Embayment and Mediterranean. van Buchem et al. present a basin scale study, constrained by Sr isotope stratigraphy and a revised biostratigraphic zonation, for the Dezful Embayment (SW Iran) that contains the prolific OligoMiocene Asmari Formation reservoirs. This paper suggests that glacio-eustatically driven sea level fluctuations strongly controlled the depositional geometries of the third-order sequences and led to the isolation of the basin and deposition of extensive evaporites during two well documented periods. The paper by Janson et al. describes a third-order, transgressive– regressive depositional sequence from the early Miocene (Burdigalian) in southern Turkey. The geometrical evolution and faunal content of the parasequence is described with reference to a 1.5 km outcrop section that provides a continuous view of the platform to basinal facies transition. This study shows a clear change from an oligotrophic fauna to a mesotrophic fauna, which the authors attribute to the significant climatic change from colder to warmer conditions that is thought to have occurred midway through the Burdigalian. In summary, the above case studies provide material on which to base stratigraphic reference

MESOZOIC AND CENOZOIC CARBONATE SYSTEMS OF THE MEDITERRANEAN

models such as the Jurassic oolitic ramp/low angle rimmed shelf sequence, the Cretaceous rudistrimmed platform/intra-shelf basin sequences, and the Oligo-Miocene coral-foraminifera dominated buildup sequences. Notably the studies that are based on continuous exposures of entire carbonate sequences along platform/ramp to basin transitions, provide the ideal geometrical framework for the integration of different analyses and are the best basis for reference models. It is also in this type of outcrop that progress can be made in our, as yet, poor understanding of the exact interaction between the biological processes and the changing physical conditions of the Earth.

Diagenetic reference models The classification of diagenetic reference models used in this volume, is based upon the geometrical relationship between the diagenetic bodies and the original sequence geometries. This creates a classification based on whether geobodies are concordant (stratabound), partially concordant (stratabound/ non-stratabound) or discordant (non-stratabound) with the original stratal surfaces. This classification implies that the original stratigraphic architecture of the target sequence is known. The expression of these three types of reference models is principally controlled by the interaction of the stratigraphic architecture of the primary sedimentary system, with the fractures and faults created during postdepositional tectonics and the timing and composition of fluid flow. A major challenge in diagenetic studies is the three-dimensional representation of diagenetic bodies. The stratabound diagenetic patterns are easiest to address because of their close link with the geometries of the sequence stratigraphic framework (e.g. Budd et al. 1995; Moore 2001; Ehrenberg et al. 2006; Lucia 2007). In rocks affected by burial diagenesis this geometrical aspect is much more difficult to document, and therefore relatively poorly illustrated in the literature. High resolution 3D seismic data and subsequent detailed seismic attribute analyses can generate images of diagenetic bodies in sequences where the acoustic impedance contrast between the host rock and the diagenetic products is sufficiently large. There are a number of examples of non-stratabound ‘porosity aureoles’ caused by hydrothermal diagenesis, which have been mapped using 3D seismic data from other parts of the world (Kidston et al. 2005). A good example of stratabound diagenesis is provided by Weidlich, who compares the effect of diagenesis at sub-aerial exposure surfaces on Triassic sequence boundaries in different climatic zones. This outcrop-based study clearly demonstrates the

5

importance of the consideration of global factors such as climate in the interpretation of diagenetic observations. Rosales & Perez-Garcı´a in an outcrop study from the Lower Cretaceous of Spain, provide a good example of a mixed stratabound/nonstratabound case study, where a detailed stratigraphic understanding is combined with a thorough diagenetic analysis. Distinction is made between sub-aerial exposure related meteoric diagenesis along the main sequence boundaries, controlled by the contrasting diagenetic responses of the primary sediment composition of the different platform environments, and later burial diagenesis caused by fluid circulation along restricted fault and fracture zones. The latter is interpreted to have had a minimal effect on the small-scale porosity and permeability of the samples studied. Another excellent example of a mixed stratabound/non-stratabound case study is presented by Sharp et al. This case study is exceptional because of the scale and continuity of the outcrop exposure upon which the study is based. The authors describe the mapping of threedimensional seismic scale exposures of midCretaceous carbonate sequences from the Zagros Mountains of Iran. The paper describes the extent to which the primary sedimentary facies and postdepositional fractures and faults have controlled the distribution and extent of diagenesis in the subsurface. The resultant reservoir is clearly nonstratabound at the seismic scale. However, detailed analyses reveal that there is a local stratal control on the distribution of porosity and permeability at the reservoir scale. A clear example of non-stratabound, burial diagenesis is presented by Ronchi et al. who present diagenetic case studies from two different thrust belt settings that is, the Southern Alps and the Apennines. In both settings, pervasive dolomitization was observed, but the dolomitization can be traced to compositionally contrasting fluids using isotope geochemistry. The authors relate this difference to the contrast in the two tectonic settings at the time of dolomite formation. In the example from the southern Alps, high sub-aerial relief generated a large hydrodynamic head that resulted in an abundant fresh water intake into the carbonate sequence. This contrasts with the tectonic setting of the Apennine example where the sequences were submarine at the time of dolomite precipitation. Although different in composition, both cases display similar diagenetic trends. The authors propose that the ‘diagenetic front’ of dolomitizing basin fluids followed the lateral migration of the developing foreland basins in each case and hence the penetration of these diagenetic fronts was ultimately controlled by the structural evolution of the individual mountain chains.

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Conclusions This volume contains a number of stratigraphic and diagenetic case studies, some elements of which may serve as the basis for age and architecture specific reference models of carbonate sequences. The development of such models is considered to be an essential step in the incremental improvement of our understanding of carbonate sequences and the controls on their development. The challenge to Earth scientists studying carbonate sequences now and in the future will be to integrate the ever-expanding volume and range of geological, geophysical, geochemical and reservoir performance information effectively to generate internally consistent models for carbonate deposition, sequence development and reservoir performance. Ideally such models, suitably scaled, will be equally applicable to academic studies, the exploration and development phases of the field life cycle and in the prediction of future reservoir performance. This Special Publication arises from the AAPG European Region Meeting ‘Reference models of Mesozoic and Cenozoic carbonate systems in Europe and the Middle East’ that was organized from 30 April to 1 May 2006, in Palma de Mallorca, Spain. We gratefully acknowledge the financial support for this publication from SHELL and a private donation by Dr M. C. M. Kabel and Mrs M. M. Kabel-Panhuysen from Warmond, the Netherlands. We wish to publicly acknowledge the time and dedication of a number of reviewers whose help has been essential in bringing this volume to print: M. Bachmann, P. Bassant, J. Bernaus, P. de Boer, P. Burgess, G. Davies, P. Farzadi, H. Hillgartner, P. Homewood, A. Horbury, P.-H. Larsen, B. Levell, J. Marcello, F. Nader, L. Pomar, M. Poppelreiter, K. Rugweid, M. Simmons, R. Swennen, B. Vincent, P. Wagner and G. Warrlich.

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shallow-water carbonates in northern Oman. Sedimentology, 49, 555– 581. P OMAR , L. 2001. Types of carbonate platforms: a genetic approach. Basin Research, 13, 313– 334. P OMAR , L. & H ALLOCK , P. 2008. Carbonate factories: a conundrum in sedimentary geology. Earth Science Reviews, 87, 134– 169. P OMAR , L. & K ENDALL , C. G. S. C. 2008. Architecture of carbonate platforms: a response to hydrodynamics and evolving ecology. In: L UKASIK , J. & S IMO , J. A. (eds) Controls on Carbonate Platform and Reef Development. SEPM Special Publication, 89, 187 –216. P URSER , B., T UCKER , M. E. & Z ENGER , D. (eds) 1994. Dolomites: A Volume in Honour of Dolomieu. International Association of Sedimentologists, Special Publication, 21. R EAD , F. 1985. Carbonate platform facies models. AAPG Bulletin, 69, 1 –21. S CHLAGER , W. 1991. Depositional bias and environmental change – important factors in sequence stratigraphy. Sedimentary Geology, 70, 109– 130. S CHLAGER , W. 2003. Benthic carbonate factories of the Phanerozoic. International Journal of Earth Science (Geologische Rundschau), 92, 445–464. S HARLAND , P. R., A RCHER , R. ET AL . 2001. Arabian Plate Sequence Stratigraphy. GeoArabia Special Publication, 2. S IMMONS , M. D., S HARLAND , P. R., C ASEY , D. M., D AVIES , R. B. & S UTCLIFFE , O. E. 2007. Arabian Plate sequence stratigraphy: potential implications for global chronostratigraphy. GeoArabia, 12, 101– 130. T ISSOT , B. 1979. Effect on prolific petroleum source rocks and major coal deposits caused by sea level changes. Nature, 277, 462–465. VAN B UCHEM , F. S. P., H UC , A. Y., P RADIER , B. & S TEFANI , M. 2005. Stratigraphic patterns in carbonate source-rock distribution: second-order to fourth-order control and sediment flux. In: H ARRIS , N. B. (ed.) The Deposition of Organic-Carbon-Rich Sediments: Models, Mechanisms and Consequences. SEPM Special Publication, 82, 191– 224. VAN B UCHEM , F. S. P., A L -H USSEINI , M., M AURER , F. & D ROSTE , H. 2010. Aptian Stratigraphy and Petroleum Habitat of the Eastern Arabian Plate. GeoArabia Special Publication, 4. W EISSERT , H. & E RBA , E. 2004. Volcanism, CO2 and palaeoclimate: a late Jurassic-early Cretaceous carbon and oxygen isotope record. Journal of the Geological Society, London, 161, 695 –702. W ISSLER , L., W EISSERT , H., M ASSE , J.-P. & B ULOT , L. G. 2002. Chemostratigraphic correlation of Barremian and lower Aptian ammonite zones and magnetic reversals. International Journal of Earth Science (Geologische Rundschau), 91, 272–279.

The influence of basin architecture and eustacy on the evolution of Tethyan Mesozoic and Cenozoic carbonate sequences K. D. GERDES1*, P. WINEFIELD2, M. D. SIMMONS3 & C. VAN OOSTERHOUT1 1

Shell International Exploration and Production bv, Kessler Park 1, 2288 GS Rijswijk, The Netherlands

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Shell International Exploration and Production, P.O. Box 481, Houston, TX 77001-0481, USA

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Neftex Petroleum Consultants Ltd, 115B Milton Park, Abingdon, OX14 4SA, UK *Corresponding author (e-mail: [email protected]) Abstract: During the Mesozoic and Cenozoic Eras, regional tectonic processes, eustatic variations and the volume and distribution of non-carbonate sediment controlled the progressive expansion and rapid reduction of the accommodation space available for the deposition of carbonate sequences, in the area that is now the Mediterranean and Middle East. We present a simplified super-regional tectonostratigraphic history of this area from earliest Triassic time to the present day, to demonstrate the influence of these large-scale processes on the evolution of major Tethyan Mesozoic and Cenozoic carbonate sequences. The time period is divided into 11 tectonostratigraphic phases (TSP) two of which (1 and 11) are incomplete. Each TSP commenced with major changes in basin architecture in response to regional tectonic processes. Subsequent pulses of transgression and regression generated sequence stratigraphic hierarchies. These stratigraphic hierarchies reflect the interaction between regional and local tectonics, eustatic variations, carbonate growth processes, climate and non-carbonate sediment supply. A map is presented of a major second-order transgressive sequence (TST) within each TSP to illustrate the maximum extent of marine onlap. These maps also include the main plate configurations; active regional tectonic features and the resultant time averaged carbonate gross depositional systems that developed during the transgression. The sequence of maps illustrate that the volume of available accommodation space during the Mesozoic and Cenozoic Eras reached a maximum during the Late Cretaceous and has been progressively reduced during the Cenozoic Era to the present day minimum.

The aim of this paper is to demonstrate the variation in accommodation space in the Tethyan region during the Mesozoic and Cenozoic Eras and the effect of this variation on the evolution of Tethyan carbonate sequences at a super-regional scale. A simplified super-regional tectonostratigraphic framework for the Mediterranean and Middle East area (Fig. 1) is presented to facilitate the link between regional tectonic and eustatic processes and the large-scale evolution of carbonate sequences (Fig. 2). This framework subdivides the time period into a set of tectonostratigraphic phases (TSPs). A sequence of maps illustrating the approximate maximum extent of carbonate deposition during a major second-order transgressive sequence (TST) within each TSP is presented to demonstrate the evolution of Tethyan carbonate sequences at a super-regional scale.

Simplified tectonostratigraphic framework We subdivide the Mesozoic and Cenozoic Eras into 11 tectonostratigraphic phases, two of which (TSP 1

and 11) are incomplete (after Sharland et al. 2001, 2004; Simmons et al. 2007). Each tectonostratigraphic phase is bounded by two super-regional sequence boundaries identified by the presence of a first-order unconformity or its correlative conformity surface. These unconformity surfaces can be correlated with plate-scale tectonic events that control significant changes in regional basin architectures. The re-orientation and fragmentation of landmasses during these tectonic events also influenced carbonate deposition by modifying oceanic circulation patterns. The hierarchies of stratigraphic sequences that developed during each TSP were controlled by the interaction of accelerated tectonic subsidence, eustatic variation, climate and carbonate growth processes. The base of TSP 1, the Permo-Triassic boundary, is defined by the scope of this review. Using this classification, it is a time boundary within the TSP, its base occurring at the Pangaean break-up unconformity (Ruban et al. 2007). The top of TSP 11 is defined as the present day and is considered to be incomplete. The durations of the TSPs have

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 9–41. DOI: 10.1144/SP329.2 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Present day setting of the Mediterranean and Middle East.

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been approximated using the timescale of Gradstein et al. (2004) and calibrated against the sequence stratigraphy of Sharland et al. (2001, 2004) as undertaken by Simmons et al. (2007). Partial TSP 1: Permo-Triassic boundary (Tr10 SB) to mid-Triassic unconformity (Tr60 SB); 22 Ma TSP 2: mid-Triassic unconformity (Tr60 SB) to earliest Jurassic unconformity (J10 SB); 28 Ma TSP 3: earliest Jurassic unconformity (J10 SB) to late Toarcian unconformity (J20 SB); 20 Ma TSP 4: late Toarcian unconformity (J20 SB) to intra-Oxfordian unconformity (J60 SB); 29 Ma TSP 5: intra-Oxfordian unconformity (J60 SB) to near-base Cretaceous unconformity (K40 SB); 15 Ma TSP 6: near-base Cretaceous unconformity (K40 SB) to late Aptian unconformity (K90 SB); 27 Ma TSP 7: late Aptian unconformity (K90 SB) to mid-Turonian unconformity (K150 SB); 22 Ma TSP 8: mid-Turonian unconformity (K150 SB) to near-base Cenozoic unconformity (Pg10 SB); 25 Ma TSP 9: near-base Cenozoic unconformity (Pg10 SB) to Oligocene unconformity (Pg30 SB); 33 Ma TSP 10: Oligocene unconformity (Pg30 SB) to Messinian unconformity (M1 SB); 28 Ma Partial TSP 11: Messinian unconformity (M1 SB) to present day; 6 Ma. The aim of this tectonostratigraphic subdivision is to provide a simplified framework that facilitates the correlation of regional scale tectonic and eustatic processes to the extent and characteristics of the respective carbonate sequences. In attempting to provide a simple framework a number of generalizations have been made. This scheme cannot attempt to address accurately the diachroneity and detail of tectonic events and facies boundaries that occur across a region of this size. This scheme is also inappropriate to describe the stratigraphic architecture in areas where local microplate motions or rapid tectonically-driven vertical movements introduce locally significant unconformities. Finally, a number of important processes relating to the tectonic evolution of the Tethyan area, such as the exact age and nature of the opening of the eastern Mediterranean, remain uncertain (Garfunkel 2005; Gardosh & Druckman 2006). For further detail on the tectonics and stratigraphy of the area the reader is referred to Dercourt et al. (2000); Sharland et al. (2001, 2004); Haq & Al-Qhatani (2005); Stampfli et al. (2001) and Ziegler (1990).

Regional controls on accommodation space and carbonate sequence development Eustacy, tectonics and sediment supply are classically regarded as the primary controls on the

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architecture of sedimentary sequences (e.g. van Wagoner et al. 1988; Vail et al. 1991; Miall 1997). These concepts were developed for siliciclastic sequences and when combined with the variation in climate and carbonate growth processes have been found to be applicable to carbonate sequences (e.g. Handford & Loucks 1993; Coe et al. 2003; Schlager 1993, 1999, 2003, 2005). A number of workers (Vail et al. 1977; Haq et al. 1987; Sharland et al. 2001; Simmons et al. 2007) have argued that eustatically-controlled marine flooding events are synchronous worldwide and thus the architecture of each sequence in a basin will be controlled by its accommodation space (a function of subsidence, eustacy and original basin bathymetry) interacting with climate and local sediment supply. In this paper we propose that initial basin bathymetry, eustacy and regional tectonics are the primary controls on accommodation space and oceanic circulation. In combination with the evolution of carbonate growth processes, climate and the distribution of non-carbonate sediments these are the main factors that control the architecture of carbonate sequences on a regional scale. Periods of rapid global climate change and significant ‘greenhouse’ periods are thought to have occurred during the Mesozoic and Cenozoic Eras (Jenkyns 2003). The periodic expulsion of large volumes of greenhouse gases during volcanic activity associated with large igneous provinces (LIP), and the subsequent oxidation of these gases, have been proposed as potential mechanisms to trigger both these climatic changes and large-scale modifications in seawater chemistry. Carbonate deposition has been proposed as a significant process in the subsequent geological sequestration of carbon in the marine system by which lower levels of carbon dioxide in the biosphere are restored. Advances in cyclostratigraphy and chemostratigraphy have led to correlations being proposed between major periods of LIP activity, orbital climate signatures, oceanic anoxic events (OAE), carbonate productivity and mass extinctions (Jenkyns 2003; Saunders & Reichow 2009). An analysis of the potential impact of these events on the global carbon budget and carbonate productivity during the Mesozoic and Cenozoic Eras is beyond the scope of this contribution. However, a number of such events are thought to have occurred during these eras and the best documented are identified in the text and Figure 2. Figures 3 and 4 illustrate how the underlying tectonically-derived relief and basin floor topography, once inherited as bathymetry, can control the extent and continuity of the accommodation space during a transgression. A transgression across a highly structured unconformity, such as that created during the syn-rift phase of extensional

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Fig. 2. Tectonostratigraphic subdivision of the Mesozoic and Cenozoic Eras used in this paper. A selection of the major regional tectonic events associated with each tectonostratigraphic phase is listed adjacent to the lowermost sequence boundaries [(SB, sequence boundary (in red); MFS, Maximum Flooding Surface (in blue); OAE, Oceanic Anoxic Event (after Jenkyns 2003); LIP, Large Igneous Province (after Saunders & Reichow 2009; Haq & Al-Qhatani 2005); nomenclature after Simmons et al. 2007; geological timescale after Gradstein et al. 2004; eustatic curve after Haq (1996)]. (a) Triassic and Jurassic periods. (b) Cretaceous and Cenozoic periods.

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Fig. 2. (Continued).

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Fig. 3. Carbonate sequence development model: rugose topography. The effect on carbonate deposition of a eustatic cycle across a rugose depositional substrate. Rapid variation in relative relief can generate facies changes over small lateral distances. The effect of siliciclastic pollution is more locally constrained by structural focusing and more distal areas of the margin may be unaffected. Relative falls in eustatic level produce isolated zones of sub-aerial exposure and potential karstification.

basin development or in a compressional thrustfaulted terrane, may create a number of discrete sub-basins, partially or completely isolated by intervening highs (Fig. 3). Such sub-basins may become periodically or permanently isolated from normal marine conditions. This basin setting is suited to

Fig. 4. Carbonate sequence development model: inclined plane or ramp. The effect on carbonate deposition of a eustatic cycle across a depositional substrate resembling an inclined plane. Along strike variation is dependent upon inherited structure, differential subsidence and organic topography across the shelf. Facies changes with large lateral extent can be generated by relatively small variations in accommodation space. Siliciclastic pollution can extend across the margin relatively unhindered by structural baffles and access the deeper parts of the basin. Relative falls in eustatic level can cause sub-aerial exposure of regionally extensive surfaces.

the development of isolated or detached carbonate sequences that typically exhibit rapid lateral facies variations over relatively short distances. Structural focusing of siliciclastic, evaporitic and/or volcaniclastic sedimentation into inboard sub-basins or away from developing carbonate platforms can protect carbonate environments from conditions that might inhibit or terminate carbonate production. In contrast, a similar transgression of identical magnitude over a relatively unstructured and gently inclined substrate, such as that created in a

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post-rift or foreland basin setting, creates a more uniform basin floor topography (Fig. 4). This type of basin setting is conducive to the development of facies belts that are laterally continuous on a regional scale. In such a setting, events such as siliciclastic pollution or the lowering of eustatic level can result in regionally significant lateral shifts in facies belts that greatly influence the development and juxtaposition of reservoir and seal facies. These figures illustrate that different substrate architectures can be developed in contrasting basin types for example, extensional and foreland basins, or at different evolutionary stages within the same basin for example syn-rift and post-rift. They further illustrate that the same eustatic perturbation can generate radically different carbonate sequence architectures in tectonically distinct but adjacent basins. Post-depositional processes may further modify the sequences observed at outcrop or in the subsurface. Subsequent sediment burial, regional tectonics and basin evolution continue to affect carbonate sediments after deposition as they control: (a) the degree of subsequent burial and/or uplift; (b) the regional and local thermal and pressure regimes; (c) the orientation and intensity of fracturing and stress; (d) the juxtaposition and nature of subsequent sequences; (e) the fluid flow at both local and regional scales. These additional factors control the nature and pervasiveness of processes such as diagenesis and fracturing which may modify the primary depositional lithologies extensively and ultimately determine the present day hydrocarbon potential of carbonate sequences.

Carbonate sequence stratigraphies and hydrocarbon potential Detailed sequence stratigraphic descriptions are beyond the scope of this analysis. However, it is instructive, to consider the hydrocarbon potential of the higher order sequences in Tethyan carbonate stratigraphic hierarchies. The transgressive architectures of second- and third-order TSTs often contain sequences with regionally significant source rock and/or seal potential. Following each second-order maximum flooding surface (MFS), a second-order high stand systems tract (HST) results from the combination of the slowing/cessation of basin subsidence and eustatic variation. Second- and third-order HST sequences contain regressive architectures that often host the rapid lateral expansion of carbonate sequences as carbonate platforms switch from

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predominantly aggradational to progradational depositional geometries. A number of prolific carbonate reservoirs of regional significance, were developed in the area of interest during third-order HSTs. In many well documented reservoirs fourth-order parasequence hierarchies are required to fully explain the lateral variation in reservoir quality and performance observed in hydrocarbon accumulations (Strohmenger et al. 2008).

Regional carbonate depositional environment maps (Figs 5 – 15) The description of each tectonostratigraphic phase summarizes the large scale tectonic events and basin processes active during the phase and their influence on the resultant architecture of carbonate sequence development. We present maps that combine the main tectonic elements active during each TSP with the gross carbonate depositional environment for a second-order transgressive systems tract (TST) that is taken to represent a proxy for the maximum extent of marine accommodation space during the TSP. Regional carbonate depositional environment maps of the following second-order TSTs, with sediment sequences (after Sharland et al. 2001, 2004; Simmons et al. 2007), are presented in Figures 5–15. TST 1: Permo-Triassic boundary (Tr10 SB) to midTriassic flood (Tr40 MFS) TST 2: mid-Triassic unconformity (Tr60 SB) to late Triassic flood (Tr80 MFS) TST 3: earliest Jurassic unconformity (J10 SB) to Toarcian flood (J10 MFS) TST 4: late Toarcian unconformity (J20 SB) to early Oxfordian flood (J50 MFS) TST 5: intra-Oxfordian unconformity (J60 SB) to Kimmeridgian flood (J100 MFS) TST 6: near-base Cretaceous unconformity (K40 SB) to mid-Aptian flood (K80 MFS) TST 7: late Aptian unconformity (K90 SB) to early Turonian flood (K140 MFS) TST 8: mid-Turonian unconformity (K150 SB) to Maastrichtian flood (K180 MFS) TST 9: near-base Cenozoic unconformity (Pg10 SB) to Palaeogene flood (Pg20 MFS) TST 10: Oligocene unconformity (Pg30 SB) to early Miocene flood (Ng20 MFS) TST 11: Messinian unconformity (M1 SB) to upper Messinian flood (M1 MFS). Each figure contains a list of time-equivalent formation names from the Mediterranean and Middle East. Colour coding indicates the hydrocarbon potential of each formation. The maps are

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integrations of the depositional environment maps of Dercourt et al. (2000) and Popov et al. (2004) with the tectonic maps of Ziegler (1990) to which the reader is referred to for further detail. Lithological ornament has been added to the maps where the facies have been confirmed by additional well and/ or outcrop information. Given the scale of the maps the locations of structural features are approximate. In the following text the Mediterranean and Middle East regions are referred to collectively as the area of interest. The area is subdivided geographically as follows: † Northern area of interest refers to present-day southern Europe; † Western area of interest refers to the present day Bay of Biscay through the Iberian Peninsula to Morocco; † Southern area of interest refers to the present day margins of North Africa; † Eastern area of interest refers to the present day Middle East and Arabian Gulf; † Central area of interest refers to the area occupied by the present day Mediterranean.

TSP 1: Permo-Triassic boundary (Tr10 SB) to mid-Triassic flood (Tr40 MFS); 22 Ma TSP 1 is a partial TSP as the base of the Triassic (Tr10) is not a sequence boundary of regional significance.

Tectonic setting During the Late Palaeozoic a series of major plate tectonic events had a profound impact upon the area of interest (Stampfli et al. 2001; Stampfli & Borel 2002; Stampfli et al. 2002; Scotese 2004; Torsvik & Cocks 2004). These events commenced during the Carboniferous with the collision of Gondwana and Laurasia and the initiation of the Hercynian Orogeny. The western and central areas of interest were closest to this plate collision and the regional hiatus associated with this orogeny is referred to as the Hercynian unconformity, a widespread and clearly recognizable sequence boundary throughout the area. During the mid-Permian, the Cimmerian Superterrane started to rift away from Gondwana to form the NeoTethys ocean (Sengor 1990; Stampfli 2000; Stampfli et al. 2001; Stampfli et al. 2002; Torsvik & Cocks 2004). Some authors correlate the NeoTethys break-up unconformity with the Hercynian unconformity and use the term Pangaean break-up unconformity to cover both hiatuses (Le Metour et al. 1995; Ruban et al. 2007). The diachroneity of these tectonic events resulted in a large variation in the surfaces upon which the earliest Mesozoic sediments were deposited.

Figure 5 shows the location of present day North Africa to the west of Iberia during this TSP. In the southwestern region, plate reorganization led to extensional stresses of the subsequent Atlantic orientation being applied obliquely across zones of preexisting crustal weakness. The resultant suite of transcurrent basins formed sites of restricted marine conditions and evaporitic sediments competed for marine accommodation with siliciclastic red beds and volcaniclastic sediments (Perrone 2006). It is not universally accepted that there was oceanic continuity between the present-day eastern Mediterranean and NeoTethys during this period. Evidence of Permian extension east of present-day Italy is limited but can be found as far west as southern Tunisia. Data from the Levant indicate that horst and graben topography developed along the southern margins of the opening NeoTethys Ocean (Hips & Argyelan 2007). The present day Arabian Platform and central Iran were passive margins on the opposing rift shoulders of the NeoTethys ocean (Sharland et al. 2001; Stampfli et al. 2001). NeoTethys was located at equatorial latitudes as the plates drifted northwards and in this basin the earliest Triassic sediments were deposited conformably on Permian sequences. The expulsion of greenhouse gases associated with volcanic activity in the Siberian Traps LIP may have contributed with other factors to modify both climate and seawater chemistry to create an environment that hosted both an oceanic anoxic event and a major mass extinction during this TSP (Saunders & Reichow 2009). The latter accounted for the extinction of up to 90% of marine taxa (Jenkyns 2003).

Gross depositional environment During TSP 1 the variation in timing and response to the break-up of Pangaea across the area of interest controlled the extent, age and nature of the sediments deposited. Figure 5 illustrates the depositional setting during the latter periods of TST1 that is, from the Permo-Triassic boundary (Tr10 SB) to mid-Triassic transgression (Tr40 MFS). Sediments of earliest Triassic age were not deposited over most of the central and southern area of interest. These areas were emergent during this TSP and the oldest Mesozoic sediments deposited on the unconformity surface in these areas were regressive siliciclastic sediments of Carnian to Norian age (TSP 2). In the SW, a suite of localized depocentres developed in restricted basins within which carbonate deposition was muted. In the NW and central regions, siliciclastic red beds, the Verrucano, dominated both this and the subsequent TSP (Perrone et al. 2006). These continental lithologies

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Fig. 5. TSP 1: Partial TSP Permo-Triassic boundary (Tr10 SB) to mid-Triassic unconformity (Tr60 SB) Approximate extent of carbonate depositional environments illustrated by palaeogeographic map of TST 1—Permo-Triassic boundary (Tr10 SB) to mid-Triassic flood (Tr40 MFS) (modified after Ziegler 1990; Dercourt et al. 2000; Simmons et al. 2007).

occurred both above and beneath mud dominated Muschelkalk carbonate ramps and evaporites. In the northern and northeastern area of interest conditions were more open marine and tectonic subsidence driven by active extension combined with eustacy to create highly variable volumes of accommodation space. Sedimentation was commonly continuous across the Permo-Triassic

boundary. Central and southern Europe was progressively transgressed during this TST (Cassinis et al. 1988; Massari & Neri 1997). Muddy ramps with shallow marine, restricted shelfal facies belts including thrombolites and microbial buildups were deposited. These ramp sequences attained considerable thicknesses in southern France, Cantabria and the Italian Dolomites (Bosellini & Neri 1991;

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Bosellini & Stefani 1991). In parts of the Southern Alps, isolated carbonate platforms with rapidly aggrading, high relief rimmed shelves were deposited in response to variations in tectonic subsidence rates (Stefani & Caputo 2002). The eastern margin of the Afro-Arabian Plate bordered NeoTethys and underwent progressive post rift subsidence. Broad carbonate ramps developed over much of the area. The Permo-Triassic boundary is clearly associated with major global extinctions and can be identified within the upper sections of the Khuff mixed shelf sequence over much of the Arabian Plate by a marked divergence in carbon isotope signature (Benton 2003). Carbonate sediments directly above the Permo-Triassic boundary typically lacked platform- constructing organisms and were often dominated by thrombolitic assemblages indicative of microbial activity in stressed environments (Virgone et al. 2008). Carbonate factories became re-established within less than 1 million years and widespread early Triassic carbonate ramp platforms were established in the passive margin sequences of the Khuff formation (Husain et al. 2008). The Scythian transgressions across the margins of NeoTethys can be traced as far SE as Socotra Island, where fossiliferous limestones directly overlie Proterozoic crystalline basement (Ballini et al. 2008). Along the opposing passive margin of NeoTethys, in present day Iran, kilometre-scale Kangan shallow-water carbonate platforms developed as a result of continuous deposition across the Permo-Triassic boundary (Insalaco et al. 2006).

TSP 2: mid-Triassic unconformity (Tr60 SB) to earliest Jurassic unconformity (J10 SB); 28 Ma Tectonic setting The unconformity at the base of this TSP coincides with an acceleration of extensive translational rifting which can be traced from the eastern seaboard of the United States through the western Tethyan region via the Azores Fracture and into the Caspian Sea. This translational rifting created a suite of northeasterly trending Triassic basins in the western Tethyan region. These rapidly subsiding basins opened along reactivated Hercynian crustal weaknesses (Oujidi et al. 2006). Extensive volcanism was associated with this extension and volcaniclastic sediments and pillow lavas were widely deposited in the western area of interest. Eurasia and the Afro-Arabian Plate were emergent and formed the northern and southwestern margins of the basinal area. Pre-existing structures such as the

Qatar Arch were rejuvenated along the eastern margin of the Afro-Arabian Plate. These structures strongly influenced facies development in carbonate sequences deposited in the low accommodation space on the partially transgressed passive margin of NeoTethys.

Gross depositional environment Figure 6 illustrates the depositional setting during TST 2 that is, from the mid-Triassic unconformity (Tr60 SB) to late Triassic transgression (Tr80 MFS). The combined effects of regional volcanism and the deposition of siliciclastic sediments derived from the denudation of the surrounding hinterlands limited the marine accommodation space available for carbonate deposition in the developing basins of the west and central areas of the region. Carbonate deposition was limited to the crests of submerged footwall blocks where platforms were structurally isolated from non-carbonate sediments (Fig. 3). The first Tethyan marine onlap in eastern Morocco has been dated as late Ladinian (CrasquinSoleau et al. 1997). In Mallorca, the open marine conditions introduced during the previous TSP gave way to unfossiliferous evaporitic facies interbedded with alkaline volcanics deposited in confined basins. In the NW and central area of interest, carbonate deposition also illustrated a strong structural influence. High relief rimmed platforms developed on extensional fault blocks (Bosellini 1984; Blendinger 1985). These platforms coalesced into broader ramps as tectonic activity waned at the end of this TST and during the subsequent HST. The intervening basinal areas were often sediment starved or occupied by distally derived volcanic and siliciclastic sediments. Extensive evaporitic and peritidal platforms covered most of the Periadriatic region. Conditions were open marine and pelagic to the south and east. Salt deposition and gypsiferous marls were widespread whilst important lacustrine source facies were deposited in the restricted, partially anoxic basins in present day Albania, Italy, Montenegro, Greece, Sicily and Libya (Zappaterra 1994; Peters et al. 2004). Transgression across the southern margins of the area of interest during the latter stages of TSP 2 led to the deposition of extensive evaporite sequences with volcanic interbeds along the North African margin from Morocco to Tunisia. The Arabian Platform was emergent during most of TSP 2 and carbonate deposition was similarly delayed until the latter stages of this TSP. Low subsidence rates and rejuvenated structures on the

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Fig. 6. TSP 2: mid-Triassic unconformity (Tr60 SB) to earliest Jurassic unconformity (J10 SB). Approximate extent of carbonate depositional environments illustrated by palaeogeographic map of TST 2—mid-Triassic unconformity (Tr60 SB) to Late Triassic flood (Tr80 MFS) (modified after Ziegler 1990; Dercourt et al. 2000; Simmons et al. 2007).

underlying broad passive margin combined to create a shelf with low accommodation. These conditions led to the development of both temporal and spatial high frequency lateral facies variations in the sequences deposited during this TSP. The transgressions spread across a broad peneplained surface

upon which hypersaline ramps such as the Kurrachine formation were deposited in present day Syria. In present day Iran, carbonate deposition was largely uninterrupted and is represented by the carbonate platforms of the Dashtak formation (Szabo & Kheradpir 1978).

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TSP 3: earliest Jurassic unconformity (J10 SB) to late Toarcian unconformity (J20 SB); 20 Ma Tectonic setting During TSP 3 the Central Atlantic/Tethys transtensional system became the principal locus for the break-up of Pangaea. Acceleration of the central Atlantic rifting led to major sinistral motion along the Azores Fracture Zone (AFZ) as Africa moved eastward. This motion was co-incident with the progressive opening of the western Mediterranean. This transcurrent motion was extended via Sicily into the eastern Mediterranean. Strongly diachronous extension commenced in the Hettangian in Italy (earliest TSP 3) and propagated eastward to reach the Rutbah Arch in Syria by the late Toarcian (latest TSP 3). By the end of TSP 3 seafloor spreading was established in the central Atlantic as far north as the AFZ. The Bay of Biscay seaway linked the Tethyan and Arctic Seas. Active rift basins were established in the Maghreb-Iberia region in the west of the area of interest. Volcanism at the onset of this phase, associated with the Central Atlantic Magmatic Province (CAMP) LIP, represented the peak of extrusive activity in the region from the Bay of Biscay, through Libya, Egypt and the Levant (Wilson & Guirard 1998). The Triassic –Jurassic boundary is characterized by a major mass extinction and evidence for oceanic anoxic conditions has been presented for both this boundary and the earliest Toarcian (Jenkyns 2003; Hesselbo et al. 2002; Beerling & Berner 2002). The possible correlation of both of these events with the rapid increase in volcanic activity associated with the CAMP LIP (Triassic – Jurassic boundary) and the Karoo-Ferrar LIP (early Toarcian) respectively is intriguing. In the central area of interest accelerated extensional tectonics caused widespread platform collapse and a subsequent major expansion of marine accommodation. Rifting also initiated along the southeastern margin of the Arabian Platform and the associated rift flank uplift may be a partial cause for the rejuvenation of the Arabian Platform hinterland that occurred during this TSP.

Gross depositional environment Figure 7 illustrates the depositional setting during TST 3 that is, from the earliest Jurassic unconformity (J10 SB) to the Toarcian transgression (J10 MFS). Regional extensional tectonics led to a major expansion in accommodation space, particularly in the north, central and western Mediterranean areas. On the northern and western margins siliciclastic input was subdued. The marginal basins of

the Atlantic and Western Tethys experienced accelerated subsidence and a reduced siliciclastic input. This resulted in a proportionally larger volume of the expanded accommodation being available for carbonate deposition. Siliciclastic input was largely restricted to the southern margins of the region where rift flank uplift along the Afro-Arabian Plate rejuvenated the hinterland. In the eastern area of interest, lower rates of subsidence along the passive margin combined with relatively low eustatic levels and continuous siliciclastic supply to restrict carbonate deposition to the latter stages of this TST when low energy hypersaline ramps such as the Marrat formation (Kuwait) were deposited (Rousseau et al. 2006; Griest et al. 2005). Carbonate deposition in the north and central area of interest was strongly influenced by the underlying structural architecture. Deposition on the rapidly subsiding, highly rugose, rifted topography led to the deposition of condensed sequences (Ammonitico Rosso) on submerged fault block crests. These sequences were often superseded by platforms of regional extent and kilometre-scale thicknesses that were subsequently drowned by the regional Toarcian maximum transgression. Carbonate and evaporitic deposition was dominant in many of the restricted basins and organic-rich sequences deposited in these settings are important source horizons in Italy, Albania, Greece and Libya (Zappaterra 1994). Deposition on the southern margins of the area of interest was similarly influenced by the underlying structural architecture. Platform and basin topographies were developed in the Atlassic, Maghrebian and Tunisian basins. Structural isolation made these basins more prone to hypersaline conditions. Evaporitic sequences deposited in the basins of eastern Algeria and Tunisia act as regional seals for the prolific siliciclastic TAGI and TAGS reservoirs in Algeria and Tunisia (MacGregor & Moody 1998). The final transgression of this phase is a major maximum flooding surface (J10 MFS, Sharland et al. 2001, 2004) and TSP 3 culminated in the widespread deposition of marine, sometimes anoxic, shales and limestones of Toarcian age.

TSP 4: late Toarcian unconformity (J20 SB) to intra-Oxfordian unconformity (J60 SB); 29 Ma Tectonic setting The unconformity at the base of TSP 4 can be correlated with an acceleration in Tethyan rifting and Pangaean break-up. The continued eastward translation of Africa created the Maghrebian-Ligurian Trough in the western Mediterranean and

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Fig. 7. TSP 3: Earliest Jurassic unconformity (J10 SB) to late Toarcian unconformity (J20 SB). Approximate extent of carbonate depositional environments illustrated by palaeogeographic map of TST 3—earliest Jurassic unconformity (J10 SB) to Toarcian flood (J10 MFS) (modified after Ziegler 1990; Dercourt et al. 2000; Simmons et al. 2007).

introduced seafloor spreading into the AlboranKalybide Basin eastward as far as Sicily (Ziegler et al. 2001). Farther east transtensional faulting created the Tethyan margins of Cyrenaica, western Egypt and the eastern Mediterranean. Large-scale structural inversion to the north of the region caused the temporary isolation of Tethys from the Arctic. Depositional hiatuses persisted in parts of

the central and southwestern area of interest until the Bajocian as tectonic relief associated with the horst and graben topography outpaced eustacy in the rapidly extending basins. The central and southern basins of NeoTethys continued to subside and basinal areas such as the Umbrian and Adriatic basins coalesced. The southern margin of the area of interest consisted of

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a series of rifted horst and graben with an overall regional deepening northward. The northeastern margin of the Afro-Arabian Plate was emergent at the beginning of TSP 4. During TSP 4, rifting commenced in the Shabwa basin of Yemen and the margin of the Afro-Arabian Plate extended from the Equator to 308N. The contrast in carbonate depositional environments caused by the different basin architectures

of the actively extending western and central area of interest and the passive margin/incipient compressional setting of the eastern area of interest was progressively accentuated during the Jurassic.

Gross depositional environment Figure 8 illustrates the depositional setting during TST 4 that is, from the late Toarcian unconformity

Fig. 8. TSP 4: Late Toarcian unconformity (J20 SB) to intra-Oxfordian unconformity (J60 SB). Approximate extent of carbonate depositional environments illustrated by palaeogeographic map of TST 4—Late Toarcian unconformity (J20 SB) to early Oxfordian flood (J50 MFS) (modified after Ziegler 1990; Dercourt et al. 2000; Simmons et al. 2007).

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(J20 SB) to early Oxfordian transgression (J50 MFS). The map demonstrates a clear contrast in second-order sequence morphology controlled by substrate architecture across the region. The rifted horst and graben topography in the western and central area of interest provided an underlying structural control that led to rapid lateral facies variations (Fig. 3). This contrasted with the peneplained northeastern margin of the Afro-Arabian Plate that formed a gently inclined low relief substrate with limited accommodation space. Facies development in Mallorca is fairly typical of structurally controlled deposition in the tectonically active, extensional settings in the west and central area of interest. At this location oolitic facies were deposited on the structural highs with time-equivalent condensed marls in the adjacent depocentres. Re-deposited oolites within turbiditic basinal deposits may indicate contemporaneous tectonic activity and/or strong current activity. Elsewhere in the central and northern areas of interest deposition was continuous with reefal facies fringing hypersaline environments on shallow-water platforms and condensed sequences (i.e. Ammonitico Rosso) being deposited on fully submerged highs. These platforms were separated by expanded and merged basinal areas in which cherty limestones and marls were deposited (Zappaterra 1994). The southern margin of the area of interest, from Morocco to Libya, consisted of a sequence of rift basins with an intervening graben, which deepened northwards. Extensive deep-water shales and marls were deposited in the graben whereas a variety of shallower water facies were developed on the intervening structural highs, some of which were not fully transgressed until the Callovian. This gross depositional environment, developed on a continuously subsiding passive margin, persisted until TSP 8 on certain sections of this margin, such as the Levantine Basin (May 1991; Roberts & Peace 2007). In the eastern area of interest the peneplained Arabian Platform was not transgressed until the Bajocian. Post-Bajocian sequences map out a progressive transgression across the Arabian Shelf (Al-Husseini 1997) (Fig. 4). Deposition on this shelf was characterized by low accommodation and reduced sediment volumes. Sequences formed typically low relief shallow marine ramps that evolved into aggradational flat-topped platforms of fine-grained carbonates. In Oman intra-platformal sequences thickened by an order of magnitude along strike of the passive margin (Rousseau et al. 2006). Along the southern margin of the Afro-Arabian Plate rapid transtensional subsidence led to the deposition of carbonate sequences in the Shabwa Basin which illustrate strong structural control (Fig. 3).

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TSP 5: intra-Oxfordian unconformity (J60 SB) to near-base Cretaceous unconformity (K40 SB); 15 Ma Tectonic setting The base of TSP 5 is defined by a widespread regional unconformity that can be correlated with accelerated crustal separation in western Tethys and the continuous translation of Africa eastwards. An acceleration in spreading rate in the central Atlantic has also been associated with this TSP. The northern and central areas of interest were bordered by progressive rifting in the North Atlantic and seafloor spreading to the northeast of the Levantine basin. The overall effect of these increases in extensional tectonics in the areas bordering the Tethyan region was to put the intervening emergent areas into mild compression. Many massifs in the area of interest were rejuvenated and acted as localized sediment provenance areas during this TSP. Progressive rifting along the northwestern margin of the Arabian Plate, the opening of the proto-Owen basin along the southeastern margin and rifting in the Shabwa Basin put the Arabian Platform into mild compression. Structural reactivation, halokinesis and sediment aggradation all contributed to the subdivision of this margin into a suite of intra-shelf basins that had a fundamental control on carbonate deposition. Many structures such as the Hamad Uplift, Qatar Arch and northeastern margin of Oman were sub-aerially exposed. However, increases in eustatic sea level, led to an expansion of accommodation space within the intervening basins. The interaction of these processes subdivided the marine accommodation space on the margin into isolated basins with restricted circulation, separated by structurally controlled highs.

Gross depositional environment Figure 9 illustrates the depositional setting during TST 5 that is, from the Oxfordian unconformity (J60 SB) to the Kimmeridgian transgression (J100 MFS). TST 5 is one of the most regionally correlatable transgressions having been identified along the eastern edge of the Arabian Shelf (Rousseau et al. 2005), in the Interior Fars province of Iran (Golestaneh 1965) and throughout southern Europe and North Africa (Atrops & Benest 1984; Zappaterra 1994). A Tithonian– Kimmeridgian OAE has been proposed by some authors (Haq & Al-Qhatani 2005). In west and central Tethys the expansion of accommodation space accelerated. The subsiding, tectonically active central platforms and the northern and southern margins of Tethys were

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Fig. 9. TSP 5: Intra-Oxfordian unconformity (J60 SB) to near-base Cretaceous unconformity (K40 SB). Approximate extent of carbonate depositional environments illustrated by palaeogeographic map of TST 5—intra-Oxfordian unconformity (J60 SB) to Kimmeridgian flood (J100 MFS) (modified after Ziegler 1990; Dercourt et al. 2000; Simmons et al. 2007).

transgressed and carbonate deposition was largely uninterrupted. Regional tectonic subsidence combined with an overall increase in eustatic signature led to the extensive deposition of ramps and low angle rimmed shelves. Large carbonate platforms developed in the central and western area of interest. Evidence of rapid alternations of sub-aerial exposure and complete submergence on platform surfaces indicate tectonic instability and/or eustatic

fluctuations during this time (Benito et al. 2005; Badenas & Aurell 2001). The southern margin of the area of interest entered a post-rift phase. Variations in bathymetry and regional subsidence were subdued and condensed carbonate sequences were deposited on submerged horsts. Thick sequences of fine-grained and bioclastic limestones were deposited along the margin. The Saharan Platform hinterland was

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rejuvenated and deltaic and shallow-marine siliciclastic sediments prograded northwards across the shelf (Ben-Ferjani et al. 1990). The combination of structural reactivation, halokinesis and sediment aggradation maintained an intra-shelf division along the Arabian Platform margin. Many re-activated structures were not transgressed until the Tithonian (Rousseau et al. 2005). Deposition on the Arabian Shelf was characterized by relatively low accommodation space and sediment volumes. Extensive carbonate ramps developed in this setting with rapid lateral facies variations (e.g. Arab, Jubaila and Hanifa formations) (Hughes et al. 2008). During the high-order lowstand and transgressive systems tracts of these carbonate sequences organic rich facies accumulated in both the marine basin fringing the Arabian Platform and within the intra-shelf basins. These sequences include hydrocarbon source horizons of regional importance. Conversely, barrier shoals deposited in the corresponding high-order high stand systems tracts are regionally significant reservoir facies (Strohmenger et al. 2004).

TSP 6: near-base Cretaceous unconformity (K40 SB) to late Aptian unconformity (K90 SB); 27 Ma Tectonic setting The unconformity at the base of TSP 6 correlates with another period of major plate reorganization. Seafloor spreading rapidly expanded in the central Atlantic and Indian Ocean. Extension accelerated in the North Atlantic and the Azores and Anatolian Fracture Zones were juxtaposed, if not connected, through the eastern Mediterranean. The resultant northeasterly drift of the AfroArabian Plate towards the southern margin of Eurasia led to sinistral transtensional stresses affecting much of the Tethyan region. Extensive transtensional rifting and subsidence, often with associated volcanics, occurred along the southern margin of Tethys from the Saharan Atlas of Morocco to the Levant. TSP 6 includes the early Aptian OAE, or ‘Selli Event’, which is one of the best documented of the proposed global OAEs (Jenkyns 2003). The duration of this OAE is broadly co-incident with increased volcanic activity in the High Arctic, Whitsunday and Ontong Java large igneous provinces (LIPs) (Saunders & Reichow 2009). Much of the hinterland bordering the region was rejuvenated. Large areas of Europe were emergent and major siliciclastic systems entered the marginal basins in the north of the area of interest. The

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Afro-Arabian hinterland along the southern margin was also rejuvenated. This may have been related to the initiation of extensive rifting along the Central African Rift Zone, which extended from Nigeria to Ethiopia. Along the eastern margins of the region the exhumation of NeoTethys commenced. The Owen Basin opened along the southeastern margin and the Arabian Platform was bordered by passive margins on three sides.

Gross depositional environment Figure 10 illustrates the depositional setting during TST 6 that is, from the near-base Cretaceous unconformity (K40 SB) to the mid-Aptian transgression (K80 MFS). Carbonate deposition clearly reflected the differences in basin evolution between eastern and western Tethys during this period of rapidly increasing eustatic sea level rise. Large carbonate platforms developed in the region in response to the expanding accommodation space. The nature of the platforms responded to local variations in substrate and rates of differential subsidence generated by the contrasting tectonic processes active across the region. Intermittent but significant pulses of siliciclastic supply led to the deposition of major deltas and periodic localized interruptions in carbonate deposition along the Arabian Shelf (Davies et al. 2002). The base of this TSP is defined by a major unconformity or hiatus that has been identified in basins from Morocco to Yemen. Deposition in the western area of interest was typified by the succession in Mallorca where the pre-existing basins were submerged beneath post rift pelagic marls of Barremian –Aptian age. In central and eastern Tethys, shallow-water carbonate sequences developed on a suite of broad rifted platforms (Apulia, Karst, Abruzzi, Lucania) that were separated by graben with time-equivalent basinal facies. Accommodation along the southern margin of the area of interest was controlled by tectonically-driven relative relief inherited from the pre-existing extensional basin architecture. This relief was superimposed upon the general northward deepening of the margin and led to the development of a sequence of mixed shelves consisting of locallyderived siliciclastic systems that periodically prograded across rugose carbonate shelves (Chaabani & Razgallah 2006). TSP 6 was a period of spectacular growth of distally steepened carbonate ramps and rimmed shelves on the Arabian Platform. Increasing tectonic subsidence on three sides of the plate combined with a series of major transgressions that expanded marine accommodation space progressively. The rejuvenated hinterland shed significant volumes of siliciclastic sediment onto the shelf and major

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Fig. 10. TSP 6: near-base Cretaceous unconformity (K40 SB) to late Aptian unconformity (K90 SB). Approximate extent of carbonate depositional environments illustrated by palaeogeographic map of TST 6—near-base Cretaceous unconformity (K40 SB) to mid-Aptian flood (K80 MFS); (modified after Ziegler 1990; Dercourt et al. 2000; Simmons et al. 2007).

delta systems, such as the Zubair, were deposited across Iraq and Kuwait (Davies et al. 2002). The Arabian Shelf was subdivided by preexisting topography into a series of intra-shelf basins and highs upon which extensive carbonate platforms developed. The cyclical aggradational and progradational deposition of rudist shoals of the Shu’aiba formation was typical of carbonate sedimentation in these settings during TSP 6. The Shu’aiba shelves formed kilometre-wide belts that fringed the Bab intrashelf basin. Organic rich, fine-

grained carbonates were contemporaneously deposited in the basin (van Buchem et al. 2002a; Yose et al. 2008; Kerans 2007). Seismic and well data illustrate a highly differentiated topography in the intra-shelf basins with slope margin dips ranging from 0.5 degrees to more than 30 degrees and water depths varying from several tens of metres to one hundred metres (Droste pers. comm. 2007). Margin progradation distances of the order of 50 km have been attributed to parasequences in third-order high stand system tracts (van Buchem

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et al. 2008; Pierson et al. 2008). Variations in lithologies, accumulation rates and rates of compaction created intra-shelf topography that further modified depositional facies. Rapid lateral facies variations developed reservoir, source and seal potential in higher order sequences.

TSP 7: late Aptian unconformity (K90 SB) to mid-Turonian unconformity (K150 SB); 22 Ma Tectonic setting The late Aptian unconformity at the base of TSP 7 coincides with the Cretaceous Magnetic Quiet Zone in the central Atlantic and is evident throughout North Africa and Europe where it can be correlated with the unconformity known as the ‘Austrian Event’. To the north and west of Tethys, seafloor spreading commenced in the North Atlantic and the Bay of Biscay separated Iberia from Europe. The Iberian Block and other localized massifs were elevated and became provenance areas for siliciclastic sediments. Seafloor spreading accelerated in western Tethys and the Alboran block was emergent and shed sediment south towards the northern margin of the Afro-Arabian Plate. North and eastern Europe were also emergent at the onset of TSP 7 and extension in the Pindos, Vardar and Zanda basins led to the emergence of the Carpathian, Hellenide and Pontide ranges along the northeastern margins of the area of interest. Siliciclastics were shed into the basins that formed along these margins. Regional plate motions combined to accelerate the northeasterly drift of the Afro-Arabian Plate towards Eurasia. Sinistral translation, with a northeasterly strike, initiated major transpressional inversion and transtensional rifting with associated volcanism along the southern margin of Tethys. The opening of the central and southern Atlantic accelerated at this time whilst in the eastern area of interest subduction of NeoTethys was locked. These regional events would have placed the Afro-Arabian Plate under compression and the hinterland of the southern margin of the area of interest was rejuvenated. This TSP also includes two well-documented OAEs that is, the early Albian ‘Paquier Event’ and the ‘Bonarelli Event’ associated with the Cenomanian –Turonian boundary (Jenkyns 2003). Both of these events can be correlated with periods of increased volcanic activity in global LIPs (Fig. 2).

Gross depositional environment Figure 11 illustrates the time-averaged depositional setting during TST 7 that is, from the late Aptian

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unconformity (K90 SB) to the early Turonian transgression (K140 MFS). During this TSP temporal and spatial variations in basin evolution resulted in large variations in carbonate deposition at both regional and local scales. Eustatic rises in sea level ensured that accommodation space expanded and carbonate deposition was prolific. The late Aptian unconformity at the base of this TSP is a clearly imaged incision surface on the intra-shelf platforms of the Arabian Platform (van Buchem et al. 2008; Yose et al. 2008; Pierson et al. 2008). Sub-aerial exposure estimates of between 3 and 5 Ma have been attributed to this strongly karstified surface (Raven et al. 2008) that has ten metres of penetrative diagenesis associated with it in Abu Dhabi (Yose et al. 2008). During TSP 7 western Tethys was dominated by post rift subsidence. Carbonate sequences were typified by deep-water organic shales and marly limestones, some with source rock potential. This general trend was interrupted on the margins of massifs such as the Iberian Plate where mixed siliciclastic/carbonate platforms developed. Differential subsidence on post-rift carbonate platforms hundreds of kilometres in width led to the deposition of complex and laterally varying sedimentary facies. Post-rift subsidence and increased eustatic signature, however, led to marine facies dominating over littoral and continental settings at a regional scale (Garcia et al. 1993). The structured nature of the substrate along the southern margin of the area of interest led to a similar diversity in facies distribution. In Tunisia and Cyrenaica in Libya, carbonate deposition was focused in deep-water marginal basins separated from the marine shelf where substantial siliciclastic systems were deposited between exposed highs (Ben-Ferjani et al. 1990; El-Hawat & Abdulsamad 2004). In the Sirte Basin of Libya successive transgressions penetrated progressively southward in the basinal areas whereas the intervening horsts remained emergent. In the Western Desert of Egypt a semi-arid shallow marine shelf developed with periodic siliciclastic incursions, marine transgressions and hypersaline conditions that is, the Abu Roash C-G formations (Said 1990; Abdel-Khalak et al. 1989). This environment extended eastwards along this northerly deepening margin from Egypt to Syria. During TSP 7, a sequence of major transgressions combined with a regional eastward tilting of the Arabian Plate to progressively increase accommodation space on the margin. The substrate resembled an inclined plane at a regional scale, in marked contrast to the structured nature of the contemporaneous depositional surface in the central and western areas of interest. Carbonate sequences of the Arabian Platform at this time recorded

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Fig. 11. TSP 7: Late Aptian unconformity (K90 SB) to mid-Turonian unconformity (K150 SB). Approximate extent of carbonate depositional environments illustrated by palaeogeographic map of TST 7—Late Aptian unconformity (K90 SB) to early Turonian flood (K140 MFS) (modified after Ziegler 1990; Dercourt et al. 2000; Simmons et al. 2007).

the spectacular growth and periodic demise of carbonate ramps and low angle rimmed shelf sequences. Periodic run-off from rejuvenated hinterlands fed major siliciclastic systems such as the Burgan Delta (Davies et al. 2002). Several large

restricted intra-shelf basins developed during TSP 7, the largest of these, the Shilaif Basin, was over 300 km wide. Differential sedimentation rates were sufficient to create intra-shelf topography within the basins during high-order TSTs and

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HSTs. A sequence of shallow water, rudist dominated, aggrading ramps and rimmed shelves (Nahr Umr, Mauddud, Natih & Mishrif formations) prograded into these basins in which time-equivalent organic rich shales were deposited (Burchette 1993; Alsharhan 1995; van Buchem et al. 2002b). In present day Iran, the initial transgression of this TSP deposited the organic rich shales of the Kazhdumi Formation within a wide intra-shelf basin. These shales were then superseded by rudist dominated carbonate ramp and rimmed shelf sequences of the Lower Sarvak and lowermost Upper Sarvak formations (Sharp et al. 2002; Razin et al. 2010). Prolific carbonate and siliciclastic reservoir sequences of regional importance were deposited during TSP 7. TSP 7 also includes the Cenomanian and Turonian periods that were times of enhanced organic matter productivity on a global scale and the high order sequences in TSP 7 include many world-class source rocks.

TSP 8: mid-Turonian unconformity (K 150 SB) to near-base Cenozoic unconformity (Pg 10 SB); 25 Ma Tectonic setting TSP 8 commenced with the initial stages of the collision between the Afro-Arabian and Eurasian plates in the centre and west of the area of interest and the closure of NeoTethys in the east. A major acceleration in North Atlantic seafloor spreading and the progressive restriction of the Bay of Biscay seaway also occurred in the west of the region (Ziegler 1990; Segura et al. 2006). Fold belts initiated along the southwestern margin of Tethys in the Atlas and Tellian regions. Along the northeastern margins of the area of interest the westward extension of the orogenic belt comprising the Carpathians, Pontides, Hellenides, Anatolides and eastern Alps created a semi- continuous barrier of fold belts and flysch basins that further isolated the Tethyan basins. The structurally enclosed Tethyan basins underwent post-rift subsidence throughout this phase. These regional plate movements initiated major diachronous tectonic events. The southern Tethyan margin underwent dextral compression during the late Santonian and this deformation, with a predominantly northeasterly strike, continued until the Oligocene on parts of this margin. This event, widely known as the ‘Syrian Arc’ tectonic event, propagated eastward from northeast Africa to Syria. Fault re-activation, folding and structural inversion occurred throughout the Tethyan marginal basins. Farther east along this margin, complementary transtensional stresses created the rapidly subsiding

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basins of the Euphrates, Sinjar and Anah Grabens in present day Levant and Syria (Guirard 1998). The progressive obduction of the Semail Ophiolite onto the Arabian Platform also occurred during TSP 8. This margin was now in compression and the associated broad regional upwarping caused a mid-Turonian unconformity of regional significance. The along strike subdivision of the Arabian Plate shelf into intra-shelf platforms and basins was still evident but the influence of this pre-existing depositional grain was over-printed by the crustal flexure associated with ophiolite obduction. A change in the regional tectonic setting is also apparent from the northwesterly orientation of many of the structures formed at this time. In the east of the area of interest the strongly asymmetric Zagros foredeep basin formed west of the Zagros Suture (Alavi 1994).

Gross depositional environment TSP 8 was characterized by global maxima in eustatic sea levels, major tectonic inversions and long periods of non-deposition (Haq et al. 1988). The complex interaction between eustacy and tectonics generated widely differing depositional surfaces across the area of interest and the carbonate sequences deposited during TSP 8 strongly reflect this variation in substrate. Figure 12 illustrates the depositional setting during TST 8 that is, from the Mid-Turonian Unconformity (K150 SB) to the Maastrichtian transgression (K180 MFS). The plate collisions and strike slip tectonics of the late Cretaceous period modified the regional setting for carbonate deposition in three main ways. The orogenic belts and ophiolite terranes along the southwestern and northeastern margins started the progressive isolation of Tethys from global oceanic circulation. The denudation of the surrounding fold belts filled the basin margins with flysch deposits. The translational tectonics across the region caused basin inversion and the elevation of large areas of the region above sea level. Carbonate deposition was gradually re-established in some areas during TSP 8 when the tectonically driven topography was subsequently submerged by transgressions that progressively reached global maxima. On the southwestern margins, sediment run-off from the Atlas, Maghrebian, Rifian and Tellian fold belts filled a reduced accommodation space with flysch and largely excluded carbonate deposition. In northwestern Tethys, strike slip basins along the Iberian margins were progressively drowned and carbonate deposition on this tectonically active substrate broadly kept pace with subsidence (Garcia et al. 1996; Segura et al. 2006). Flysch derived from the semi-continuous orogenic chains that fringed the central and northern

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Fig. 12. TSP 8: Mid-Turonian unconformity (K150 SB) to near-base Cenozoic unconformity (Pg10 SB). Approximate extent of carbonate depositional environments illustrated by palaeogeographic map of TST 8—Mid-Turonian unconformity (K150 SB) to Maastrichtian flood (K180 MFS) (modified after Ziegler 1990; Dercourt et al. 2000; Simmons et al. 2007).

margin of Tethys filled many marginal basins and carbonate deposition was limited to the deeper, more distal basins. In the central area of interest the rapid expansion of accommodation space led to high levels of

carbonate production on the Apulian, Dinaric and Adria platforms (Zappaterra 1994). Large volumes of sediment were deposited as carbonate turbidites in the intervening Umbria-Marche pelagic basins (Montanari et al. 1989; Zappaterra 1994; Argnani

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et al. 2006). By the end of this TSP the central Tethyan platforms had converged onto the southern margin of Eurasia. The timing of this convergence has been dated as Campanian from the onset of flysch deposition (Marroni et al. 1992). Along the southern margin of the area of interest, and east of the Tellian-Atlas foldbelts, carbonate deposition was controlled by rejuvenated basement structures and newly formed features created by active translational ‘Syrian Arc’ tectonics. The sequence of TSP 8 transgressions progressively submerged the flanks of the Tunisian platform, the Sirte Basin graben complex, the Abu Gharadig, Sinjar, Euphrates and Anah Grabens. In the Sirte basin and Western Desert of Egypt mixed carbonate and siliciclastic sequences were deposited on shallowwater platforms whilst organic rich shales and carbonates were deposited in the grabens (Said 1990; Keeley & Massoud 1998). Submerged Syrian Arc trend structures may have formed important submarine ramps that accentuated restriction and enhanced source rock potential in the Sirte Graben and the Gulf of Suez (Sirte Shale and Brown Limestone). Farther east along the margin carbonate deposition extended south of the present day Gulf of Suez. Senonian inversion structures in Cyrenaica, the Levant and Syria were eventually submerged by the end of TSP 8 (Anketell 1996; Roberts & Peace 2007; Caron & Mouty 2007). The final transgression associated with TSP 8 is correlated with a global eustatic maximum and deposited extensive coccolith-rich sediments over much of the area of interest. The Arabian Platform underwent significant inversion at the onset of TSP 8. Structural highs were repeatedly rejuvenated and the scale of these inversions largely negated the effect of eustatic increases. Evidence of exposure and condensed sedimentation during the early stages of TSP 8 is common in areas such as Oman and Iran (van Buchem pers. comm.). The mid-Turonian unconformity at the base of TSP 8 removed hundreds of meters of section from the northeastern margin of the Arabian Plate. This unconformity is regionally correlatable as the surface between the Wasia and Aruma groups; Mishrif and Khasib; Sarvak and Surgah; Mishrif and Laffan; Natih and Fiqa formations at different locations across the margin (Sharland et al. 2001). Much of the platform was emergent and flysch from the bordering ophiolite belts was deposited on the margin from the Campanian onwards. Carbonate deposition was of limited extent in this area for much of the early part of this TSP (Filbrandt et al. 2006). Some of the earliest areas transgressed were relict synclines overlying previous intra-shelf basins in central Iraq (Sharland et al. 2001). These initial transgressions deposited condensed sequences of glauconitic and evaporitic

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shallow-water sediments. The southwestern margin of the platform underwent continuous flexural subsidence and the most westerly areas of the margin in Saudi Arabia and Yemen were submerged during the Campanian.

TSP 9 Near-base Cenozoic unconformity (Pg10 SB) to Oligocene unconformity (Pg 30 SB) 33 Ma Tectonic setting The major plate reorganizations that occurred at the Cretaceous/Cenozoic boundary caused a series of tectonic events that dramatically reduced accommodation space. The continued convergence of the Afro-Arabian Plate with the southern margin of Eurasia accelerated the reduction of accommodation space available for carbonate deposition. This contrasts with the preceding Mesozoic Era, which was characterized by an almost continuous expansion of accommodation space in the area of interest. Within the Tethyan basin, the Apulian carbonate platform collided with Eurasia thereby initiating the western Alps. On the northwestern Tethyan margin Mesozoic normal faults were inverted as Iberia rotated clockwise and western Iberia became emergent. Plate collisions in the Betic, Pyrenean and Sardinian regions progressively closed the Bay of Biscay seaway. The combination of Afro-Arabian convergence on Eurasia and Tethyan seafloor spreading in the Alboran basin led to the obduction of ophiolites in North Africa. At the end of the Cretaceous period, ophiolite obduction waned and the resultant crustal unloading caused the emergence of most of the Arabian Platform, with the exception of a NW–SE trending foreland basin in western Iran. This margin remained emergent throughout the Paleocene and was not transgressed until the Eocene. The southern margin of the Arabian Plate acted as a clastic sediment provenance area in the Yemen. Throughout TSP 9 the Afro-Arabian Plate margin was positioned between the Equator and 258N. Volcanic activity that occurred between 65 –55 Ma ago associated with the Deccan Traps LIP (Hoffman et al. 2000) is thought to have contributed to the Paleocene-Eocene Thermal Maximum (PETM) at the Paleocene-Eocene boundary. The PETM is one of the most documented warming events in Earth’s history and the oxidation and subsequent oceanic absorption of excess carbon is thought to have caused a rapid shallowing of the calcite compensation depth (CCD) of more than two kilometres at this boundary (Zachos et al. 2005). This event is thought to have contributed to the significant palaeontological and geochemical

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changes that have been observed at this boundary and which characterized both the volume and nature of subsequent carbonate deposition (Jenkyns 2003).

Gross depositional environment Figure 13 illustrates the gross depositional setting during TST 9 that is, from the near-base Cenozoic

Unconformity (Pg10 SB) to the Palaeogene transgression (Pg20 MFS). The progressive reduction in accommodation space caused carbonate deposition to effectively cease in the area that is now central and western Europe. The expanding regions of compressional tectonics and continental convergence shed sediments into the remaining marine accommodation space and carbonate deposition

Fig. 13. TSP 9: Near-base Cenozoic unconformity (Pg10 SB) to Oligocene unconformity (Pg30 SB). Approximate extent of carbonate depositional environments illustrated by palaeogeographic map of TST 9—near-base Cenozoic unconformity (Pg10 SB) to Palaeogene flood (Pg20 MFS) (modified after Ziegler 1990; Dercourt et al. 2000; Simmons et al. 2007).

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was confined to the margins of greatly reduced basinal areas. The convergence of the Afro-Arabian and Eurasian plates further isolated Tethys from global oceanic circulation and enhanced levels of organic matter accumulated in the basinal areas. The southern and western margins of Tethys and the distal margin of the Arabian Plate became the sites of foredeep troughs and foreland basins that filled with flysch shed from the orogenic belts in the north. The main area of continuous deposition occurred along the southern margins of Tethys in the foredeep basins of the emerging Alps between the Algerian Atlas and Egypt. The marine substrate morphology was controlled by crustal flexure (Fig. 4). Carbonate deposition on these inclined planes generated ramp to low angle rimmed shelf sequences. Low-angle carbonate shelf sequences formed broad shallowwater platforms fringed with mounded nummulitoid banks upon which restricted, hypersaline conditions periodically developed. Coralgal bioherms colonized the shallower platform margins whilst organic rich shales with source rock potential developed in the deeper ramp settings. The nummulitoid reservoirs of the El Garia and Gialo formations and the Intisar bioherms of the Sirte basin were deposited during TSP 9. The final phase of ophiolite obduction along the northeastern margin of the Arabian Platform caused a widespread depositional hiatus during TSP 9. It has been estimated that over 500 metres of section were eroded from areas in northern Iraq and Syria. Carbonate deposition on the Arabian Platform was re-established in the Eocene along the central and southeastern platform margins. Subsidence was driven by tectonic loading and created an easterly dipping ramp substrate with limited accommodation space. Extensive carbonate low-angle rimmed shelf sequences with restricted platform environments, as illustrated by the Umm-er-Radhuma and Dammam formations, were deposited over much of the Arabian Platform.

TSP 10 Oligocene unconformity (Pg30 SB) to Messinian unconformity (M1 SB) – 28 Ma Tectonic setting The reduction in accommodation space available for carbonate deposition continued to accelerate during TSP 10 and sites appropriate for the nucleation of carbonate organisms diminished in both number and scale. Regional tectonics remained predominantly compressional, which led to a further reduction in accommodation space and the progressive isolation of the Tethyan basins. In western

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Tethys the rapid translation of Iberia and adjacent microplates led to the closures of the Bay of Biscay seaway and the link to the Atlantic via the Gibraltar Straits. The opening of the Valencia, Ligurian and Provencal Basins forced the Alboran and Kabylide blocks to separate and migrate southwards onto the Afro-Arabian Plate. In North Africa, the Atlas Trough was inverted and rejuvenated hinterlands surrounded the remaining marine areas. The isolation of Tethys in the north was completed. Additional Mesozoic platforms from the north and central Tethyan basin were absorbed into the Alpine orogeny. Inversion along translational fault systems on the northern margin of Tethys led to rapid unroofing that has persisted to the Present Day (Massironi et al. 2006). The southern Tethyan margin became subdivided into a series of isolated shelves and basins in which hypersaline conditions developed periodically. In the east of the area of interest, the Arabian Plate has remained largely emergent from the beginning of TSP 10 until the present day. The closure of NeoTethys was completed and subsidence continued to accelerate in the foredeep basins along the eastern margin (Sharland et al. 2001). Continent-continent collision led to the initiation of the rapidly emerging Zagros thrust front. Regional plate movements that Increased accommodation space were largely confined to the sinistral movement along the Dead Sea transform that initiated the separation of Arabia from Africa and created a suite of connected extensional marine basins in the Red Sea.

Gross depositional environment Figure 14 illustrates the depositional setting during TST 10 that is, from the Oligocene Unconformity (Pg30 SB) to the early Miocene transgression (Ng20 MFS). The siliciclastic supply from the northern Tethyan margin was sufficient to prevent carbonate deposition in the Apennine and Alpine Foredeeps (Ricci Lucchi 1986; Mattioni et al. 2006). Tethys was now a marine area of limited extent surrounded by emergent hinterlands that shed large volumes of siliciclastic debris into the basin. Carbonate deposition where present, was characterized by restriction and hypersalinity. Major climatic changes are also thought to have further inhibited the extent of carbonate deposition (Zachos et al. 1993, 2001). The limited basinal areas were highly structured and the carbonate sequences deposited were strongly influenced by substrate morphology (Fig. 3). Coralgal bioherms formed medium, to high angle rimmed shelves on the submerged footwalls and crests of thrust slices in compressional terranes in areas such as Cyprus, Sicily, Malta, Calabria and the

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Fig. 14. TSP 10: Oligocene unconformity (Pg30 SB) to Messinian unconformity (M1 SB). Approximate extent of carbonate depositional environments illustrated by palaeogeographic map of TST 10—Oligocene unconformity (Pg30 SB) to early Miocene flood (Ng20 MFS) (modified after Ziegler 1990; Popov et al. 2004; Simmons et al. 2007).

Balearics in the western Tethys basin. Similar bioherms colonized tilted fault blocks in the extensional basins of the Valencia Trough, Gulf of Suez and Red Sea. Structural isolation, glacio-eustatic sea level fluctuations (Abreu & Haddad 1998) and arid to semi-arid climatic conditions led to frequent periods of desiccation and anhydrite deposition

in the basins and the development of restricted environments on small structurally discrete platforms (e.g. Asmari formation in Zagros; van Buchem et al. 2010). In the eastern area of interest restricted carbonate facies were deposited on a hypersaline shelf of limited extent. Marginal marine early Miocene

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sequences have been correlated along the margin of the Arabian Gulf from Oman to southwestern Iran (Jones & Racey 1994; Dill et al. 2005; Dill & Henjes-Kunst 2007).

TSP 11 Messinian unconformity (M1 SB) to present day (6 Ma) TSP11 is incomplete as the TSP ends at the present day.

Tectonic setting The base of TSP 11 marks the final collision between Africa and Europe and this event completed the isolation of the Mediterranean. The western Mediterranean was isolated from the Atlantic by the closure of the Gibraltar Straits and a continuous fold belt across the Sicily Channel separated the Eastern and Western Mediterranean basins until the Plio-Pleistocene. Translational movement occurred along the Gulf of Aqaba/Dead Sea rift and in the Valencia Trough where Minorca was translated eastward relative to Mallorca. Regional compression led to minor inversions along reactivated structural trends on the southern Tethyan margin. In the eastern area of interest the Zagros foredeep also became structurally isolated. Areas of significant accommodation space were limited to the Red Sea and the Gulfs of Suez, Aqaba and Aden. Phases of tectonism and transgression persist to the present day and have caused further modification to the regional setting since the Pliocene unconformity. The eastern and western Mediterranean basins were eventually re-connected by the collision of the North African margin with the Sardinia-Corsica block which led to back-arc extension in the Tyrrhenian Sea and the rotation southeastward of Calabria. A complex suite of rapidly subsiding basins controlled by major shear zones facilitated a marine connection in the vicinity of Sicily. Open marine conditions were re-introduced to the entire Mediterranean, the Zagros Foredeep and the Gulf of Suez/Red Sea basins during the latter stages of this TSP. The connection between the eastern Mediterranean and the Gulf of Suez/ Red Sea basin was finally closed during the Pleistocene.

Gross depositional environment Figure 15 illustrates the depositional setting at the end of TST 11 that is, from the Messinian Unconformity (M1 SB) to the Upper Messinian Flood (M1 MFS). Regional tectonics and the lowering of eustatic levels reduced accommodation space to the lowest levels since the onset of the Mesozoic Era. The area available for carbonate deposition

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was restricted to a suite of structurally isolated basins within which small, laterally impersistent carbonate platforms developed. Carbonate deposition was restricted to the margins of enclosed deep hypersaline basins. The ‘Messinian Salinity Crisis’ affected the Mediterranean during this TST and led to the widespread deposition of evaporite deposits, up to kilometres in thickness, in the basin centres. Coralgal reef deposition was restricted to the basin margins. Reefs nucleated on thrust slices in compressive terranes (central and eastern Mediterranean) or on tilted fault blocks in extensional basins (Valencia Trough, Gulf of Suez, Red Sea). Marine conditions were also restricted in the Zagros Foredeep where progressive isolation led to evaporitic deposition and sabkha formation on the basin margins. Stressed carbonate environments persisted throughout Pleistocene times. Continuous flank uplift during the Recent further reduced marine accommodation space and resulted in the comparatively limited areas of present-day carbonate deposition in the Mediterranean and Middle East regions (Fig. 1).

Summary During the Mesozoic and Cenozoic Eras, regional tectonics and eustatic processes controlled the expansion and subsequent reduction of accommodation space in the present-day Mediterranean and Middle East regions. As the Mesozoic Era progressed, the volume of accommodation space rapidly expanded, primarily through a combination of crustal extension and eustacy, and reached a peak during the Late Cretaceous. The carbonate sequences developed during this Era were strongly influenced by the evolution and expansion of carbonate growth processes, eustacy, climate and the underlying basin architecture and substrate morphologies. The increase in carbonate productivity induced by the evolution and expansion of carbonate-growth processes enabled deposition to continue during periods when regional tectonics rejuvenated the margin hinterlands and siliciclastic deposition was also significant. Plate reorganization periodically isolated large marine areas from oceanic circulation and this resulted in the development of conditions conducive to the deposition of rich and extensive carbonate source rocks. The progressive destruction of accommodation space commenced in the Late Cretaceous and persists to the present day. The main controls on this process were regional compressional tectonics driven by plate collision and the lowering of eustatic levels. The accommodation space available for carbonate deposition was further reduced by

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Fig. 15. TSP 11: partial TSP: Messinian unconformity (M1 SB) to present day. Approximate extent of carbonate depositional environments illustrated by palaeogeographic map of TST 11—Messinian unconformity (M1 SB) to Upper Messinian flood (M1 MFS) (modified after Ziegler 1990; Popov et al. 2004; Simmons et al. 2007).

the increased run-off of siliciclastic sediment derived from the denudation of the emergent orogenic belts bordering the Tethyan basins. The area of present day carbonate deposition is relatively

minor when compared to the areal extent of the carbonate sequences deposited during the periods of prolific carbonate productivity of the Mesozoic Era.

REGIONAL CONTROLS ON CARBONATE SEQUENCES The authors are grateful to Shell International E and P and Neftex Petroleum Consultants for permission to publish this paper. The final presentation of the regional maps is the result of the diligent work of Ferdinand Bleijswijk, Peter Boer and Kate Wall to whom the authors are extremely grateful. Jim Markello, Peter Sharland, Bruce Levell and Gerard Stampfli are thanked for their helpful reviews of the original manuscript.

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Depositional sequences and palynology of Triassic carbonate-evaporite platform deposits in the Palmyrides, Syria ˙ 3, G. FORSˇEK1, D. TAKACˇ3, J. BUBNIC˙3 & G. KOCH4 D. LUCˇIC´1,2*, Z. IVKOVIC 1

INA, Industrija nafte, d.d., Naftaplin, Exploration Department, Branch Office, Damascus, P.O. Box 33392, Syria 2

Present address: RWE Dea NA/ME GmbH-Libya Branch, Barcelona Str., 15/A Hai Andalus, Tripoli, Libya

3

INA, Industrija nafte, d.d., Naftaplin, Exploration Division, Subiceva 29, Zagreb, Croatia 4

Croatian Geological Survey, Sachsova 2, Zagreb, Croatia

*Corresponding author (e-mail: [email protected]) Abstract: This paper describes the results of new stratigraphic correlations and petroleum systems analyses of the Triassic deposits in the Syrian Palmyrides. The correlations are based on well data and new palynological work which reconcile previous lithological and stratigraphic miscorrelations of Triassic sequences from this region. The sequences are subdivided into four Megacycles, which are directly related to the key elements of the petroleum systems of the Palmyrides. The hydrocarbon discoveries within these systems, the fluid contacts and hydrodynamics are explained with reference to the Megacycle subdivision. This subdivision is applied on a regional scale and is used to define the distribution of reservoir seal pairs for volumetric estimations.

The Triassic evaporite/carbonate successions and their associated lithofacies in the Palmyrides have recently become increasingly attractive exploration targets. Notable hydrocarbon discoveries in these facies include the oil and gas/condensate fields in the central part of the Northern Palmyrides; Cheriffe, Ash Shaer and Bilas etc, discovered by Marathon Oil Co., in the early eighties. In the late nineties, the hydrocarbon exploration in this area was continued by INA Oil Co. Commercial quantities of gas/condensate and oil have been discovered in the Kurrachine Dolomite Formation; D1 and C2 reservoirs (Jihar, Jazal and Al Mahr fields). Detailed studies of the stratigraphy and petrography of carbonate-evaporite facies completed during this exploration work programme provided new insights into local carbonate/evaporate patterns within the Hayan Block in the Palmyrides (Fig. 1). Additionally, new reserves of hydrocarbons have been discovered deeper in the Middle and Lower Triassic section within the KDF D2/2 reservoir and Amanus Shale Formation reservoirs (Jihar, Jazal and Mazrur fields). Two periods of evaporite deposition resulted in the complex and variable appearance of salt deposits and caused lateral variations in reservoir property distribution. In addition to the tectonic setting, traditional lithostratigraphical nomenclature and the previous imprecise dating of Triassic deposits further complicated stratigraphical correlation. A recently obtained well database (2D/ 3D seismic, logs, well test results, cores, cuttings)

allowed a new approach to stratigraphic correlation and enabled a seal/reservoir pair definition to be assigned. New methods in defining evaporitecarbonate distribution through geological time have been applied and the deposition has been defined in terms of megacycles. Megacycle boundaries (low impedance salt bodies) can be traced on seismic lines allowing more confident mapping and prospect delineation. Maximum flooding surfaces (MFS) have been defined using the scheme established in Sharland et al. (2001). The sequence stratigraphic scheme and interpretation of regional MFS presented in this paper is considered as a preliminary framework and awaits further confirmation by future detailed petrographical, sedimentological and biostratigraphical analyses and interpretation.

Regional tectonic setting of the NE Arabian Plate from Triassic period to Recent Syrian geology includes elements from the Precambrian (the initial stage, structuring and formation of the Gondwana continent) to the present Arabian Plate. The position and the orientation of the plate was influenced by plate tectonic movements, which led to changes in the prevailing climatic conditions, depositional environments and tectonic-structural settings. Various authors

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 43– 63. DOI: 10.1144/SP329.3 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Location map of studied area on the Arabian Plate.

DEPOSITIONAL SEQUENCES AND PALYNOLOGY

classified these tectonic and geological events using different criteria (Argyriadis 1980; Robertson & Dixon 1984; Robertson et al. 1991; Dercourt et al. 1986; Beydoun 1991; May 1991; Walley 1998; Stampfli et al. 2001; Scotese 2001; Ziegler 2001). Recently, Sharland et al. (2001), defined five main evolutionary phases for the Phanerozoic of the Arabian Plate, which are subdivided into eleven tectonostratigraphic megasequences TMS (AP1– AP11). Numerous local and semi-regional tectonostratigraphic cycles may also be recognized. The present day location of the study area lies in the Palmyrides and on the northern margin of the Arabian Plate in the vicinity of the active zone of collision with Eurasia. The present day distribution of tectonic provinces and structural boundaries on the northern margin of the Arabian Plate, was controlled mainly by Cenozoic tectonic events. The Palmyrides are considered as a type example of an intracontinental transpressive mountain belt and represent the most significant structural chain in central Syria (Ponikarov 1966a, b, 1967a, b; Al-Saad et al. 1992; Chaimov et al. 1992, 1993; Barazangi et al. 1993; Brew et al. 2001). The Palmyrides strike N458E from the Anti-Lebanon mountain chain and the Dead Sea fault system to the Euphrates Graben in the NE, beneath which they plunge. The Palmyride mountain chain is 400 km long and 100 km wide with a maximum elevation of 1300 m. The mountain belt is flanked

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by two structural highs, the Aleppo Plateau at NW and Rutbah Uplift towards SE. Geophysical anomalies below and around the Palmyrides can be explained by the existence of a suture/shear zone which was active during the Palaeozoic, Mesozoic and part of the Cenozoic. A tectonically active depocentre, the Palmyra Trough, containing 6000 m of Palaeozoic and 5000 m of Mesozoic and Cenozoic sediments formed within this shear zone. Starting in the Cretaceous and continuing throughout the Eocene and Neogene (Chaimov et al. 1992), these sediments were uplifted and inverted into their current position. The area is still under compression. Inversion processes included folding, reverse faulting, translation and small block rotation along numerous strike-slip faults (Fig. 2). The Palmyrides can be divided into northern and southern tectonic domains, by the approximately east –west oriented Jihar fault and overlying Al Daww depression (McBride et al. 1990; Searle 1994). The oldest rocks in the study area, which crop out in the cores of Hayan and Mazar anticlines in Southern Palmyrides, are Upper Triassic marls and gypsum (Lucˇic´ 2001). Jurassic deposits comprise dolomite and limestone with minor shale. Lower Cretaceous deposits are represented by sandstones and shales. Pre-Coniacian sediments have been deposited on a shallow carbonate platform as a thick dolomite-limestone succession which is sporadically intercalated with layers of shale and

Fig. 2. Simplified kinematics of the Palmyrides showing stress orientation and related strike slip tectonics, modified after Chaimov et al. 1993.

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anhydrite (Lucˇic´ et al. 2002). From the Santonian to the Oligocene, the depositional setting is represented by a deeper basin within which predominantly marls and marly limestones have been deposited. Miocene deposits consist of sandstones, conglomerates, gypsum and marls. Quaternary sediments cover almost all depressions and valleys (Ponikarov 1966a, b).

Triassic deposits on the NE Arabian Plate During Triassic time, following the disintegration of Gondwana in the Permian, Cimeria broke away northwards from Gondwana, creating the NeoTethys Ocean. At the same time, the Apulia Plate began to move northwards away from Africa along the major dextral transform fault forming the western edge of the Arabian Plate and opening up the eastern Mediterranean (Fig. 3a, b). These plate movements resulted in the passive margin that now forms the present day Palmyrides changing orientation from eastward to westward facing. (Best et al. 1993). At the onset of eastern Mediterranean opening, post-rift subsidence and transgression continued towards the east of Arabia (in the Palmyrides, Brew et al. 2001). As the Arabian Plate moved into lower latitudes the combination of a warmer climate and a widespread marine transgression introduced carbonate deposition onto the northern Pangaean shelf during the latest Palaeozoic and subsequent Triassic times. The similar depositional conditions and lithofacies distribution in this period, bounded by the ‘break-up’ unconformity of Late Permian age and the strong rifting in Mediterranean during the Early Jurassic, permit the correlation of Triassic deposits across the wider area of the Arabian Plate and Levant. (Metwalli et al. 1974; Sharief 1982, 1983; May 1991; Sadooni 1995; Ziegler 2001;

Sadooni & Alsharhan 2004). Sequence stratigraphic maximum flooding surfaces can be traced almost all over the entire Arabian Plate at that time (Sharland et al. 2001). This time period correlates with the MFS Tr10 to Tr60 and intra J10 of Sharland et al. (2001). In Early Triassic times, the Arabian Plate persisted as a relatively peneplained ENE-sloping passive margin platform. The facies pattern around the Arabian Plate indicates deposition of shallow marine silty shales, gypsiferous shales and shallowmarine carbonates. In the Levant region, Jordan and Syria transgressive-regressive cycles formed an alternating sequence of shallow-marine sands, shales and marly limestones. The deposits comprise the Sudair Formation (Saudi Arabia, UAE, Oman), Mahil Formation (Oman), Amanus Shale Formation (Syria), Beduh and Mirga Mir Formations (Iraq) and their regional equivalents from MFS Tr10 to intra Tr40. During the Middle Triassic in Saudi Arabia, deposition took place in a continental to shoreline setting in which alluvial fans from the western uplands encroached over coastal lagoons onto a platform with restricted marine conditions. Numerous evaporite salinas occupied an area between Kuwait and the southern Gulf (Ziegler 2001). Evaporites grade southward into dolomitic mudstone and fine-grained clastic sediments. In the Zagros foreland, carbonates consist of shallowing-upward cycles. The area north of the Arabian Arch was generally poor in siliciclastics. Platform carbonates were deposited on the Levant side of the plate, whereas the northern margin is rimmed by clastic-evaporitic coastal to deltaic deposits. In the Palmyra Trough shallow-marine shales, carbonates and evaporites accumulated locally. Sediments of the Jilh Formation (Arabian Peninsula), Gulailah (eastern Gulf ), Geli Khana (Iraq), lower part of

Fig. 3. (a) Paleogeographic reconstruction north west Tethys during Triassic time, after Scotese, 2001. (b) Paleogeographic reconstruction and facies distribution during Late Triassic in Syria, after Brew et al. 2001.

DEPOSITIONAL SEQUENCES AND PALYNOLOGY

Kurrachine Dolomite Formation (Syria) correlate with MFS Tr40 to intra Tr60 of Sharland et al. (2001). In the Late Triassic, across the Gulf region, clastic deposits were continuously derived from the southern edge of the Arabian Shield. The Zagros region was occupied by shallow-marine carbonates and evaporites. During the second phase of NeoTethyan extension, in the eastern part of the Arabian Plate, the drowning of the northeastern margin and local volcanic activity took place. The northern and western parts of the plate were exposed and a wide shallow-marine carbonate shelf surrounded the exposed shield. During the Carnian salinity crisis, large amounts of lowstand evaporites accumulated in the Palmyra and Sinjar troughs. This time period spanned the deposition of the Minjur (Arabian Peninsula), Kurrachine Dolomite, Kurrachine Anhydrite, Butmah, Adayah, Allan, Muss, Sergelu (Syria), Kurra Chine (Iraq), Dashtak (Iran) formations and their regional equivalents (MFS Tr60 to intra J10) of Sharland et al. (2001).

confusing nomenclatures have been used, and definite age determinations of the formations were not available. In this lithostratigraphical hierarchy, the Triassic deposits in the Palmyrides are divided into; Lower Triassic Amanus Shale Formation (Scytian), and the Middle–Upper Triassic Kurrachine Dolomite Formation (Anisian–Ladinian–Carnian). The Upper Triassic is divided into the Kurrachine Anhydrite (Carnian), Butmah, Adayah, Muss, Allan, (Norian) and Sergelu (Rethian) formations. The Mulussa group (A –E) was established by Marathon and comprised of the all Triassic and Jurassic carbonate deposits. The Kurrachine Dolomite Formation was further subdivided into informal units based on reservoir properties (Chouquette et al. 1994; Craig et al. 1994; Blanchard et al. 1995; Lucˇic´ et al. 2003a; INA & Marathon Reports). The classification introduced by INA and the Syrian Petroleum Company, SPC) is now used in Syria (Fig. 4): (1) (2)

Triassic deposits in Syria Generally, in Syria, Triassic deposits are divided into Early Triassic syn-rift and Middle to Late Triassic post-rift sediments. During the Early to Middle Triassic, clastic deposition was replaced with predominantly carbonate shelf sediments with decreasing shale content and subordinate evaporites. In the Late Triassic, much of present day Syria was part of a wide restricted, shallow shelf. Sea level fluctuations, climate changes and tectonically influenced palaeorelief controlled changes in bathymetry and shelf geometry. Carbonate deposits are intercalated with shale and thick evaporite; in places salt layers are more than 800 m thick. Similar depositional conditions prevailed during the Carnian along the entire western margin of the NeoTethys (Fig. 3a, b). These Triassic deposits are comparable with the Alpine Triassic (Druckman et al. 1982). However, in several deep offshore wells in the Adriatic Sea, the thick portion of Carnian salt is isochronous with the salt defined in Syrian wells. (Lucˇic´ 2003b).

Triassic sequences in the Palmyrides region of Syria The Triassic stratigraphy of the Palmyrides was previously based on well data (Fig. 4) by Bebeshev et al. (1988), McBride et al. (1990), Best et al. (1993), Yaroshenko & Bach Imam (1995), Searle (1994), Jamal et al. (2000), Brew et al. (2001), Lucˇic´ et al. (2003a). Different, and sometimes

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(3) (4)

The ‘E’ member was assigned to the Amanus Shale formation; The ‘D2’ member was divided as pre-salt ‘D2/2’, ‘D2 salt’ and post-salt ‘D2/1’ Kurrachine Dolomite member; The previous ‘D’ member became the ‘D1’ Kurrachine Dolomite member; The ‘C’ member was subdivided between the ‘C2’ Kurrachine Dolomite member and the Kurrachine Anhydrite Formation (previous ‘C1’ member).

Subsequent dating of the Kurrachine Dolomite Formation as Middle Triassic (Anisian–Ladinian) in age caused inconsistencies between local and regional correlations (Bebeshev et al. 1988; Yaroshenko & Imam 1995; Syrian Petroleum Co. and Marathon Reports). The most recent results by INA indicate that the Kurrachine Dolomite Formation in D2/2 below salt is of Middle Triassic (Anisian –Ladinian) age whereas the Kurrachine Dolomite Formation post D2 salt that is, the D1 and the entire C2, belong to the Upper Triassic (Carnian). The Triassic deposits of the north and east Arabian Plate were assigned to the AP6 Megasequence by Sharland et al. (2001). These authors identified eight Triassic maximum flooding surfaces. A preliminary correlation based on well logs and cuttings data can be made between the sequences observed in the Central Palmyrides and the classification of Sharland et al. (2001). More detailed biostratigraphic and sedimentological analyses are required to confirm precisely this correlation. Based on the authors’ observations, maximum flooding surfaces from the Lower Triassic (Tr10 –Tr30), Middle Triassic (Tr40 –Tr50)

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Fig. 4. Comparison between lithostratigraphic units and megacycles.

DEPOSITIONAL SEQUENCES AND PALYNOLOGY

and Upper Triassic (Tr60 –Tr80) have been identified in the Central Palmyrides wells and correlated with wells in Iraq and Iran (Fig. 5).

Triassic megacycles Results from deep wells in the central Palmyrides (Hayan Block), recently drilled by INA Oil Co., have led to a modification of the conclusions of previous studies. In Lucˇic´ et al. (2004), the ‘megacycle’ concept has been used to simplify the stratigraphy of the carbonate-evaporite cycles in the Central Palmyrides. So far four megacycles have been identified (Fig. 6). Megacycles probably coincide with second-order sequences which can be correlated on seismic profiles (Fig. 7). The spatial distribution of megacycles within the studied area is shown in Figure 8.

‘Megacycle A’: Late Permian – Middle Triassic (Scythian – Anisian/Ladinian), Amanus Shale Fm-Kurrachine Dolomite Fm (D2/2) This megacycle consists of a shallowing upward 300 m thick succession from basinal or distal ramp deposits to lower shelf– ramp sediments up to shallower intertidal –supratidal deposits. The cycle is capped with D2 salt and/or an anhydrite equivalent layer of various thicknesses up to 120 m (Fig. 9).

Lithology Silty mudstones, shales and dolomites, partly calcareous (calcitic) as well as silty and argillaceous with a few occurrences of limestones, predominate the lower part of this megacycle (Fig. 10.1-9). Medium dark grey mudstones contain silty quartz grains, pyrite and glauconite at the dolomitic base which is sporadically clayey. With the increasing dolomite content, the mudstone grades into dolomudstone and argillaceous dolomite. The shale is medium dark grey to grey, slightly dolomitic, rich in dark organic matter and pyrite, with sporadic quartz grains of silt –sand size. The silty argillaceous dolomite is greenish grey, and sporadically greyish red (Fig. 10.6), very finely crystalline to finely crystalline with mainly silty quartz grains. Dolomites are grey to greenish grey, in places anhydritized, finely crystalline to very finely crystalline with anhedral to subhedral crystals forming xenotopic to hypidiotopic fabrics. Dark organic matter is present in very finely crystalline varieties. Very light grey calcareous (calcitic) dolomites occur with relicts of fossil remains (shells of molluscs and foraminifera and fragments

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of echinodermata), suggesting dolomitization of fossiliferous limestone (probably of wackestone type). Limestones are white and the dolomitic wackestone is very light grey, containing fossil remains and dark organic matter impregnations. In places, limestones are bioclastic, peloidal and selectively dolomitized.

Biostratigraphy Lower part: Amanus Shale Formation. The biostratigraphic analysis was performed on core and cuttings material. The scarce microfossil assemblage has been found to contain Ammodiscus sp., Glomospirella sp. and Ophthalmidiium sp. suggesting a Permian–Triassic? stratigraphic range. The palynoflora is represented by an undiversified assemblage of sporomorphs and a very rich and diversified assemblage of phytoplankton, for example, acritarchs and prasinophyceans. The sporomorph assemblage consists of Densoisporites nejburgii, Densoisporites spp., Endosporites papillatus, Lundbladispora sp., Lunatisporites pellucidus, Voltziaceaesporites heteromorphus and Protohaploxypinus spp. which characterize independently dated Early Triassic successions (Visscher & Brugman 1981; Warrington 1996). This type of palynoflora may well be correlated with the independently dated Early Triassic nejburgii-heteromorphus phase in the Alpine region (Brugman 1986) and the Bundsandstein palynoflora of the Germanic facies (Orłowska-Zwolin´ska 1984; Reitz 1985). In the circum-Mediterranean region, an almost identical assemblage has been recovered from the subsurface in Libya (Adloff et al. 1986), Levant (Eshet 1990) and Syria (Yaroshenko & Bash Imam 1995). Towards the eastern Tethys, similar assemblages have been found within independently dated Early Triasic successions in Australia and described as Kraeuselisporites saeptatus zone (Dolby & Balme 1976) and Protohaploxypinus samoilovichii zone (Helby et al. 1987) respectively. Within the sporomorph assemblage of the Amanus Shale Formation, several reworked Permian taxa, for example, Cyclotriletes sp., Perisaccus granulatus, Lueckisporites virkkiae, have also been recognized. Very rich phytoplankton assemblages contain diverse acritarch taxa Veryhachium spp., Michrystridium spp. as well as undifferentiated acritarchs (Conaletes sp. sensu Brugman 1986) and prasinophycean taxa Leiosphaeridia spp., Dictyotidium sp., and Tasmanites sp. According to the higher thermal rank, several acritarchs, for example, Veryhachium spp. and Michrystridium spp, have probably been reworked from the Permian successions. Although acritarchs do not have stratigraphic significance, their mass occurrences are a recognizable feature

50 ˇ IC´ ET AL. D. LUC Fig. 5. Preliminary correlation of megacycles between Palmyrides and wider area in Iraq and Iran. Tr10, represents a shale layer in the base of the Amanus Shale Formation. Tr20, is correlated with the base of a dolomite layer in the Amanus Shale. Tr30, is a dolomite layer within shales in Amanus Shale Formation. Tr40, is correlated with carbonate layer at the base of D2/2. Tr50, possibly identified as a limestone layer in the upper part of D2/2 and is thought to predate the salt. Tr60, may represent the first carbonate layer deposited after D2 salt in Kurrachine Dolomite. Tr70, is correlated with carbonate layer above evaporate succession of Kurrachine Anhydrite Formation. Tr80, Norian carbonate layer at the base of the open marine dolomite Allan/Muss Formation, following the clastic deposits of Adayah Formation.

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Fig. 6. Megacycles interpretation of Triassic deposits in Palmyrides.

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Fig. 7. Seismic expression of megacycles.

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Fig. 8. Schematic map of megacycle distribution with posted location of profile in Figure 7.

within Early Triassic successions throughout the world (Brugman 1986). According to the palynofacies features, the depositional environment is interpreted as a proximal tidal flat with mixed terrestrial (sporomorphs) and open sea influence (Verychahium, Michrystridium) in which deposition took place under dysoxic conditions in a relatively quiet setting. The very high abundance of phytoplankton and their absolute predomination over all other palynomorphs, as well as a certain portion of amorphous organic matter (AOM), may indicate maximum flooding surfaces (MFS) and transgression (Steffen & Gorin 1993; Tyson 1995; Rameil et al. 2000; Go¨tz et al. 2003). The Permian –Triassic boundary in Syria after Bebeshev et al. (1988) lies at the base of the variegated sandstones which overlie the ‘last carbonates’ of probably Late Permian age. Upper part: Kurrachine Dolomite Formation, D2/2 member. The biostratigraphic analysis was performed on core and cuttings material. The palynoflora consists of a relatively diversified sporomorph assemblage and an undiversified acritarch and prasynophycean assemblage. Sporomorphs are represented by Todisporites spp., Concavisporites spp, Kuglerina meieri, Heliosaccus dimorphus, Microcachryidites spp., Brachysaccus spp., Ovalipollis pseudoalatus, Staurosaccites quadrifidus, Partitisporites verrucosa, Partitisporites spp., Duplicisporites granulatus, Duplicisporites spp., Aulisporites astigmosus, and Camerosporites secatus. The palynofloral assemblage of this interval is characterized by the occurrence of taxa of the Circumpolles group, for example, Partitisporites, Duplicisporites, and Camerosporites, which mark

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the base of the Ladinian worldwide within the phytogeographic province of the palaeoequatorial belt where these types of pollen have been produced (Visscher & Brugman 1981; Van Der Eem 1983; Warrington 1996). The age significant taxon of the palynofloristic assemblage, which defines the Ladinian age of this interval, is represented by Heliosaccus dimorphus. An almost identical palynological composition has been recovered from type sections in the Alpine facies (Scheuring 1978; Van Der Eem 1983; Blendinger 1988) and described as secatusdimorphus phase (Van Der Eem 1983). This type of palynoflora occurs in the Lettenkohle successions of the Germanic facies (Orłowska-Zwolin´ska 1983; Adloff et al. 1984; Reitz 1985; Heunisch 1986) which may well be correlated with the Ladinian palynoflora of the Circum-Mediterranean subsurface successions, for example, in Libya (Adloff et al. 1986), Levant (Eshet 1990) and in Syria where it has been described as the Stellapollenites thiergartii-Heliosaccus dimorphus assemblage and the Camerosporites secatus-Kuglerina meieri assemblage (Yaroshenko & Bash Imam 1995). In the eastern Tethys of Australia, similar assemblages, described as Staurosaccites quadrifidus zone, have been found within successions dated as Late Anisian –Ladinian (Dolby & Balme 1976; Helby et al. 1987). The phytoplankton taxa of the acritarch and prasinophyceans occur in relatively low frequencies and are represented by Micrhystridium spp, Crassosphaera sp. and Leiosphaeridia spp. Palynofacies features suggest a depositional setting of a restricted lagoon characterized by the accumulation of predominantly amorphous organic matter (AOM) with some open sea influence (Micrhystridium). The organic matter of the deposits of this interval has been syngenetically affected by oxygen depleted conditions. The palynological organic matter is in some samples dominated by terrestrial organic matter which may indicate periodic fluviatile incursions.

‘Megacycle B’ Upper Triassic (Carnian), Kurrachine Dolomite Fm (D2/1-D1-C2)-Kurrachine Anhydrite Fm (C1) This very thick, (over 600 m) ‘megacycle’ commences with a regional transgressive, deepening event which marked the onset of shale-dominated sedimentation. The sedimentation that followed took place in a broad, shallow basin with negligible bottom relief (1 m to 10 m over distances of kilometres) and is represented by numerous rhythmically repeated cycles that began with clay/silt

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Fig. 9. Correlation of several wells within Hayan Block, flattened at top of Triassic presenting reservoir/seal pattern.

DEPOSITIONAL SEQUENCES AND PALYNOLOGY

Fig. 10. Megacycle A selected photographs.

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shales and ended with calcium sulphates. The area was restricted and repeatedly isolated from the sea. The basin finally became lacustrine, or an inland-sea system, where marine incursions were brief. The climate was dominantly arid. The end of the cycle is marked by the deposition of a very thick salt body. Salt tectonics has formed undulating salt bodies up to 800 m thick in some areas, rather than the classical diapiric salt flow. Salt thickness increases over synclinal areas and decreases over anticlinal crests.

Lithology Lower part: In the lower part of the megacycle (Kurrachine Dolomite Fm D2/1, D1) three main lithofacies are distinguished: † Micrite limestone with interbeds of dark grey calcitic-dolomitic shales. Limestones are characterized with high carbonate mud content, bioturbated mudstone to wackestone with infrequent fossil grains, mostly ostracod tests, rare echinoid bioclasts and filaments of pelagic bivalves. † Limestone-dolomite shallowing-upward cycles. Cycles commence with stylolitized subtidal mudstone limestone and in the upper part is represented by either selectively late diagenetic dolomitized and anhydritized fenestral LLH stromatolites, or recrystallized early diagenetic stromatolite dolomites with desiccation cracks (Fig. 11.8), disrupt laminae, rare nodules of anhydrite and mould voids of gypsum. † Early diagenetic dolomicrites with nodules of anhydrite and late diagenetic dolomites. Alternation of early diagenetic recrystallized dolomicrites and stromatolitic dolomicrites with anhydrite nodules with packages of macrocrystalline late diagenetic dolomites (Fig. 11.7) with secondary anhydrite. Middle part: (Kurrachine Dolomite Fm, C2) consists of rhythmical alternations of various rock types: anhydrite, claystone (dark varved dolomudstone), dolomudstone, microporous dolomite, lime mudstone/wackestone, clayey lime mudstone/ wackestone, dark grey shale (Figs 11.1, 11.4, 11.9). Each cycle is asymmetrical, beginning at its base with dark-grey to nearly black shale and terminating with a capping anhydrite. Quite often they are laminated or thin to thick bedded with ‘chicken-wire’ or nodular structure typical of supratidal/sabkha anhydrites. In places anhydrite contain dark olive-grey laminae of dolomitic claystone (dolomudstone). A few anhydrites overlie carbonates that contain stromatolites, desiccation cracks and rarely fenestrae, suggesting deposition in intertidal zones. Some

anhydrites contain laminae of dolomitic mudstone or claystone. Upper part: (Kurrachine Anhydrite Fm) contains predominately thick portions of salt (halite) (Figs 11.2-3, 11.5-6) intercalated with claystones, shales and anhydrite. The succession is mostly free of carbonates. In places, the salt intercalation is more than 800 m thick because of tectonic overprinting. The initial thickness of the salt deposit is estimated from undisturbed sections of seismic data to have been 250–350 m.

Biostratigraphy The foraminiferal assemblage of Nodosaria ordinate, Dentalina hoi, Frondicularia woodwardi, Frondicularia sp., Agathammina sp., Ammodiscus sp. Nodosaria sp. indicates a wider Triassic stratigraphic range. The palynoflora of the interval is composed of a rich and diversified assemblage of sporomorphs dominated by pollen taxa. Phytoplankton, represented by prasinophyceans, acritarchs and coenobial alga, occur in low frequencies. The co-occurrence of Minutosaccus crenulatus, Samaropollenites speciosus, Patinasporites densus, Pseudenzonalasporites summus, Froehlichsporites traversei, Partitisporites maljavkinae, Partitisporites verrucosa, Duplicisporites granulatus, and Camerosporites secatus indicates a Carnian age of the analysed samples from this megacycle. The stratigraphic ranges of these sporomorphs in Europe (Visscher & Brugman 1981; Schulz & Heunisch 2005) and assemblages that have been described from type sections of the Alpine region in northern Italy (Van Der Eem 1983; Blendinger 1988; Roghi 2004) as well as the Carnian palynostratigraphic relationships worldwide (Warrington 1996) support this dating. In particular, the described palynological assemblage may well be correlated with the Middle Carnian (Julian) palynological densusmaljavkinae phase (Van Der Eem 1983) and the upper part of the Concentricisporites bianulatus assemblage and the Duplicisporites continuus assemblage (Roghi 2004) of the type sections in the Alpine region. The dating of this interval is also supported by the presence of Sellaspora rugoverrucata and Kuglerina meieri as well as Uvaesporites gadensis, Kyrtomisporis ervei, Staurosaccites quadrifidus and Lueckisporites junior which are not known from strata younger than Middle Carnian (Julian) and Late Carnian (Tuvalian) age (Visscher & Brugman 1981; Brugman 1986), respectively. With the exception of some typical western and southern Tethyan elements (Minutosaccus crenulatus, Samaropollenites speciosus), the assemblages of the interval are similar to those from the time equivalents of the Germanic

DEPOSITIONAL SEQUENCES AND PALYNOLOGY

Fig. 11. Megacycle B selected photographs.

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Keuper facies (Scheuring 1970; OrłowskaZwolin´ska 1983; Adloff et al. 1984; Reitz 1985). The afore-mentioned assemblages also contain several taxa such as the spores Todisporites sp., Concavisporites sp., Verrucosisporites sp., Guttatisporites sp., Porcellispora longdonensis as well as pollen Enzonalasporites vigens, Accinctisporites ligatus, Ovalipollis pseudoalatus, Triadispora sp., Brachysaccus sp., Alisporites sp., Platysaccus sp., Lunatisporites acutus, Aulisporites astigmosus. In the Carnian of the circumMediterranean region, almost identical assemblages have been recognized from an independently dated section from Sicily (Visscher & Krystyn 1978; Buratti & Carrillat 2002) and from the subsurface of Libya (Adloff et al. 1986) and the Levant (Eshet 1990). In Syria, the palynofloral composition within the upper part of the Kurachine and Butmah Formations in wells Abba-1 and Habari-1 as well as within the successions that crop out in Jabel Hayan (northern and middle Syria), described as the Patinasporites densus-Samaropollenites speciosus assemblage and dated as Carnian in age by Yaroshenko & Bash Imam (1995), is practically identical with the palynoflora of the herein analysed interval. On the regional scale, Lucˇic´ et al. 2003b have shown that the palynoflora from the upper part of Kurachine Formation in the Syrian wells (evaporitic facies) is almost identical to the palynoflora of the evaporitic facies from the Croatian offshore wells indicating Middle Carnian (Julian) age, at least for the part of the successions developed mainly as a salt (halite) body within the latest Ladinian– Norian evaporitic facies. The palynoflora of this interval may well be correlated with the Carnian palynoflora of eastern Northern America (Litwin & Ash 1993) as well as with the palynoflora of the eastern Tethyan region in Australia designated as Samaropollenites speciosus zone and dated as Carnian (Dolby & Balme 1976; Helby et al. 1987) and Carnian –early Norian (Nicoll & Foster 1994), respectively. The palynoflora of the herein analysed interval (Megacycle B) contains regular occurrences of Samaropollenites speciosus, Falcisporites stabilis, Staurosaccites quadrifidus and Enzonalasporites spp. and therefore belongs to the ‘Onslow’ microflora (Dolby & Balme 1976) which reflects the vegetation that has spread across the western and southern Tethyan continental margins within a palaeoequatorial belt of low northern and southern palaeolatitudes (Visscher & Van Der Zwan 1981; Cirilli & Eshet 1991). Prasinophycens Leiosphaedridia spp. and acritachs Veryhachium spp. and Michrystridium spp. are relatively diversified but of low frequency, in particular the coenobial chlorophycean alga (Chlorococcales) Plaeisodictyon mosellanum.

Palynofacies features suggest a depositional environment of a restricted lagoon with some influence from fresh water and open sea, mostly under oxygen depleted conditions. The sedimentation took place in a broad, shallow shelf of very low relief and wide extent which for the most part was restricted and repeatedly isolated from the sea, that is, a subtidal to intertidal and supratidal environment, with sporadic sabkha or fluviatile conditions.

‘Megacycle C’ Upper Triassic (Carnian– Norian), Butmah, Adayah, Allan, Muss Fm After the Carnian salt crisis, the environment switched to restricted shelf again with carbonate, anhydrite deposition alternating with shale (Fig. 12.1-6) (Butmah Fm). Progressive subsidence resulted in the deposition of a ‘clean’ open shelf dolomite (Norian, Allan/Muss Formation). The thickness of this cycle is about 400 m.

Lithology The basal part of this megacycle is predominantly composed of dolomicrite altered with pelitic sediments and nodular anhydrite, subordinate clayey dolomite/marlstone/dolomitic mudstone, chloritized claystone/abundant shale with pyrite and organic matter. The Allan/Muss formation at the top of the megacycle is represented by clean, medium crystalline, late digenetic dolomite.

Biostratigraphy The foraminiferal assemblage containing Glomospira inconstans, Earlandinita ladinica, Ammodiscus sp., Trochammina almtalensis indicates a wide Triassic stratigraphic range. The palynoflora of this interval contains a relatively undiversified assemblage of Concavisporites spp., Samaropollenites speciousus, Minutosaccus crenulatus, Falcisporites spp., Enzonalasporites spp., Patinasporites sp., Partitisporites spp., Camerosporites secatus, and Corollina sp. On the basis of the known ranges of these sporomorphs, the co-occurrence of Camerosporites secatus and Corollina sp. suggests earliest Norian since the last appearance of the former is generally accepted as a key taxon for the top Carnian and the first appearance of the latter usually marks the earliest Norian (Visscher & Brugman 1981; Warrington 1996). Almost identical palynofloral compositions have been reported from the subsurface of Libya (Adloff et al. 1986) and the Levant (Eshet 1990). The analysed interval shows a

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Fig. 12. Megacycle C selected photographs.

palynofloral composition which is transitional between the Patinasporites densus-Samaropollenites speciosus assemblage and the Corollina meyerianaKyrtomispora corrugatus assemblage as recognized in Syria (Yaroshenko & Bash Imam 1995) suggesting earliest Norian age for this megacycle. Similar palynofloral assemblages have been described from the eastern Tethyan region in Australia as the Minutosaccus crenulatus zone and dated as Norian (Dolby & Balme 1976; Helby et al. 1987) and late Norian –early Rhaetian

(Nicoll & Foster 1994), respectively. The palynofloral assemblages of the Syrian interval is still a part of the ‘Onslow’ microflora. Phytoplankton assemblages contain few acritarch and prasinophceans and only relatively poorly preserved dinoflagellate cysts (Rhaetogonyaulax sp.) have been recognized. These palynofacies indicate that deposition occurred in a restricted shallow marine environment that was intertidal to open marine subtidal in the upper part.

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‘Megacycle D’ Upper Triassic (Norian – Rhaetian), Sergelu Fm The last, Rhaetian ‘megacycle’ followed the submergence of the Norian carbonate platform which led to the deposition of the Sergelu Fm, over shallow shelf dolomites. The cycle ended during the regional inversion of the entire area which created an extensive Triassic –Jurassic unconformity. In the western part of the central Palmyrides, the fourth ‘megacycle’ and part of the third ‘megacycle’ is missing caused by this uplift and erosion rather than non-deposition.

Lithology Early diagenetic anhydrite with stromatolite relics, dolomicrite, early diagenetic dolomite, dedolomite (calcium carbonate-cement in fractures), volcaniclastic, pyroclastic fragments predominantly crystal, vitric tuffs, quartz crystal shards, volcanic glass with hematite-pyrite and iron-oxides, tuffite/altered volcanic glass/claystone? (chloritization, hematitization of fragments), tectonized carbonates (partly hydrothermal altered?). Fossiliferous peloidal limestone selectively dolomitized-clotted pelmicrite (WS/PS), dolopelsparite (GS) with anhydrite.

Biostratigraphy The foraminiferal assemblage of Glomospira inconstans, Earlandinita ladinica and Ammodiscus sp. and Trochammina almtalensis defines a Norian – lower Rhaetian stratigraphic range. The palynofloral composition of this interval reveals a poorly diversified sporomorph assemblage of Concavisporites spp., Samaropollenites speciousus, Minutosaccus crenulatus, Enzonalasporites sp., Patinasporites sp., Falcisporites spp. and Corollina sp. Although the assemblages of this interval are similar to those previously described, Camerosporites secatus has not been identified, which may well indicate that this interval is younger than Carnian. From the palynostratigraphical point of view in the eastern Tethyan region of Australia, the last appearance of Samaropollenites speciosus marks the top of the Minutosaccus crenulatus zone which is around the Norian –Rhaetian boundary (Dolby & Balme 1976; Helby et al. 1987) and early– middle Rhaetian (Nicoll & Foster 1994) respectively. Almost identical palynofloral compositions have been reported from the subsurface of Libya (Adloff et al. 1986), Levant (Eshet 1990) and from Syria where it has been encountered within successive palynofloras of the Corollina meyeriana – Kyrtomisporites corrugatus assemblage, Corollina meyeriana–Vesicaspora schemeli assemblage and

Corollina meyeriana– Converrucosisporites cameroni assemblage (Yaroshenko & Bash Imam 1995) and dated as Norian–Rhaetian. As with the assemblages of the previous interval, Samaropollenites speciosus and other taxa indicate the affinity with the ‘Onslow’ microflora. Phytoplankton is represented by very low frequencies with some acritarchs and prasinophyceans and Rhaetogonyaulax sp. Palynofacies characteristics suggest a restricted shallow marine, intertidal to supratidal depositional environment with brief periodic open marine influence.

Petroleum systems in the Northern Palmyrides (central part) Multiple reservoir/seal pairs can be identified in the Triassic deposits in the central part of the northern Palmyrides (Fig. 9). Two separate hydrocarbon play systems were defined. The Amanus Shale Fm and Kurrachine Dolomite Fm (D2/2) reservoir consists of dolomites & limestone saturated with oil, gas and condensate. The traps are sealed by the D2 salt. The D2 salt horizon also represents the boundary between Megacycle A and B. Furthermore, it is the boundary between the Lower and Middle Triassic and in this analysis, the D2 salt separates hydrocarbon hydrodynamic units. The second hydrocarbon play system typically consists of oil-bearing carbonate reservoirs of D1 and lowermost C2 Kurrachine dolomite passing upward to uppermost C2 carbonate saturated with gas and condensate, rarely oil. The vertical hydrocarbon distribution is gravitationally driven inside each hydrodynamic unit. Well data show that reservoir saturation by certain fluids is determined by the depth of occurrence, suggesting secondary, post migration fluid spill over, most probably during the latest compressional phase of basin inversion. Reservoir porosity is a combination of mouldic and vuggy porosity, some primary fenestral and intergranular porosity enhanced by secondary fractures. Traps are sealed by evaporites of the Kurrachine Anhydrite Formation which represents the end member of Megacycle B. The most prominent Triassic source rock horizons occur in the Lower and Middle Triassic shales of the Amanus Shale and Kurrachine Dolomite D2/2 formations and Carnian C2 reservoir (Mitchels & Malartre 2007). This distribution suggests ‘intraformational’ migration from the Al Daw depression into anticlinal traps on the northern flank of the depression. Although the identification of the source rock was unambiguous, fluid

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correlation could not be proven. The observation of the same isotopic fingerprint of fluids in both hydrodynamic units suggests that the D2 salt was breached by faults during migration providing only local entrapment mechanisms.

Z. Maric-Djurekovic, B. Sokolovic (all from INA, Research Laboratory Department). The authors gratefully appreciate and acknowledge the reviews by K. D. Gerdes, F. H. Nader and K. Ruckwied, whose comments significantly improved this paper.

Conclusion

References

Until now, the Triassic deposits in the central part of the Palmyrides have only been described locally and in terms of lithostratigraphy. However, because of several important recent hydrocarbon discoveries from deep drilling, a revised stratigraphic correlation and petroleum system definition is needed. In this contribution we have provided simple megacycle interpretations supported by detailed palynological work. The Triassic deposits are characterized by the episodic and regional deposition of salt/evaporite couplets and other specific lithofacies form natural regional seal– reservoir patterns (carbonate-evaporite intercalations). Four megacycles are defined, the first two representing shallowing-upward successions of carbonate deposits capped by regional salt/evaporate deposits of laterally variable thickness. The older, D2 salt (top of Megacycle A) represents the Lower– Middle Triassic boundary and clearly separates two hydrocarbon units in the central Palmyrides. The thick ‘main’ salt at the top of second Megacycle B represents the seal for the C2 reservoir and is comparable with the regional regressive Carnian crisis and can be traced along the western palaeomargin of the Tethys Ocean. The third megacycle is the only one representing a deepening event (transgression) during the Upper Triassic. The last megacycle has been defined at the end of the Upper Triassic and is presumed to be regressive in origin because it terminates with carbonate platform emergence followed by a prominent erosional unconformity at the Triassic –Jurassic boundary. The Triassic deposits in the Palmyrides have also been correlated with equivalent deposits on the Arabian Plate using sequence stratigraphic sequence definitions. New palaeontological data (palynology) provide improved dating and assist with the regional correlation of typical MFS. However, there is a need for more detailed study of the biostratigraphic and sedimentological/petrographic studies to improve the understanding of this stratigraphic interval.

A DLOFF , M. C., A PPIA , C., D OUBINGER , J. & L IENHARDT , M.-J. 1984. Zonation palynostratigraphiques dans les se´ries triassiques traverse´e par des sondages dans le Jura et le Bas-Dauphine´. Geologie de la France, 1 –2, 3– 21. A DLOFF , M. C., D OUBINGER , J., M ASSA , D. & V ACHARD , D. 1986. Trias de Tripolitaine (Lybie). Nouvells done´es biostratigraphiques et palynologiques. Revue de l’Institut Franc¸ais du Pe´trole, 41, 27–72 A L -S AAD , D., S AWAF , T., G EBRAN , A., B ARAZANGI , M., B EST , J. A. & C HAIMOV , T. A. 1992. Crustal structure of Central Syria: the intracontinental Palmyride mountain belt. Tectonophysics, 207, 345–358. A LSDORF , D., B ARANZANGI , M., L ITAK , R., S EBER , D., S AWAF , T. & A L -S AAD , D. 1995. The intraplate Euphrates fault system – Palmyrides mountain belt junction and relationship to Arabian plate boundary tectonics. Annali di Geofisica, 38, 385–397. A RGYRIADIS , I. 1980. The opening of the Mesozoic Tethys between Eurasia and Arabia-Africa. In: A UBOUIN , J., D EBELMANS , J. & L ATREILLE , M. (eds) Geology of the Alpine Chains Born of the Tethys. Memoire du Bureau de Recherches Geologiques et Minieres, No. 115, Orleans, 199–215. B ARAZANGI , M., S EBER , D., C HAIMOV , T., B EST , J. & L ITAK , R. 1993. Tectonic evolution of the northern Arabian plate in western Syria. In: B OSCHI , E. ET AL . (eds) Recent Evolution and Seismicity of the Mediterranean Region. Kluwer Academic Publishers, Amsterdam, 117–140. B EBESHEV , I. J., D ZHAILOV , YU . M., P ORTNYAGINA , L. A., Y UDIN , G. T., M UALLA , A., Z AZA , T. & J USEF , A. 1988. Triassic stratigraphy of Syria. International Geology Review, 30, 1292– 1301. B EST , J. A., B ARAZANGI , M., A L -S AAD , D., S AWAF , T. & G EBRAN , A. 1990. Bouger gravity trends and crustal structure of the Palmyride Mountain Belt and surrounding northern Arabian platform in Syria. Geology, 18, 1235– 1239. B EST , J. A., B ARAZANGI , M., A L -S AAD , D., S AWAF , T. & G EBRAN , A. 1993. Continental margin evolution of the northern Arabian platform in Syria. AAPG Bulletin, 77/2, 173– 193. B EYDOUN , Z. R. 1991. Arabian plate hydrocarbon geology and potential plate tectonic approach. AAPG Studies in Geology. B LANCHARD , D. C., D EMBICKI , H., J R . & W ALLACE , G. D. 1995. A petroleum system in search of a reservoir; Palmyride region, Syria. Abstract, AAPG Bulletin, 79/8, 1199. B LENDINGER , E. 1988. Palynostratigraphy of the late Ladinian and Carnian in the Southern dolomites. Review of Palaeobotany and Palynology, 53, 329– 348.

The authors gratefully acknowledge the permission of INA Plc, Exploration Section management to publish this article. We also wish to thank many geoscientists who have provided unpublished photographs and analyses, particularly D. Stankovic, V. Veseli and K. Krizmanic for biostratigraphical data, petrographical analysis were provided by I. Mesic, P. Ivak, N. Raskaj, Z. Ivanicek,

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Spatial and temporal distribution of ooids along a Jurassic carbonate ramp: Amellago outcrop transect, High-Atlas, Morocco A. PIERRE1,3,4*, C. DURLET1, P. RAZIN2 & E. H. CHELLAI3 1

UMR CNRS Bioge´osciences, Universite´ de Bourgogne – 6, Boulevard Gabriel-21000, Dijon, France 2

Institut EGID, Universite´ Bordeaux III-1, alle´e Daguin, 33607 Pessac, France 3

Faculte´ des Sciences Semlalia, PO Box 2390, Marrakech, Morocco

4

Present address: Chevron Energy Technology Company, 6001, Bollinger Canyon Road, San Ramon, CA 94583, USA *Corresponding author (e-mail: [email protected]) Abstract: Carbonate ramp systems are widespread throughout the geological record, but very few areas have seismic-scale, continuous and structurally undeformed outcrops that allow reliable interpretation of facies distributions and stacking patterns. The Amellago outcrop shows the detailed depositional and stratigraphic relationships of an ooid-dominated ramp system that is almost completely exposed along a dip profile (37 km long and 1000 m thick) in the Lower to Middle Jurassic of the southern High Atlas, Morocco. Ammonite and brachiopod fauna provide excellent biostratigraphic control on small scale stacking patterns. At Amellago, the evolution of depositional environments is evident at different scales of space and time during this period of tectonic quiescence dominated by thermal subsidence. An important observation is that the Amellago ramp system contains micrite-rich, ooid-free intervals that alternate with ooid-rich intervals. The ooid-rich intervals are mainly in the late transgressive and highstand system tracts, whereas the ooid-free intervals occur in the early transgressive phase. More than 25 such alternations were recorded in high frequency cycles and at the scale of one large cycle at the Aalenian/Bajocian transition. These compositional changes and the associated different ramp geometries are interpreted to result from the combined effects of eustatic sea level and climatic changes.

Carbonate ramps display a wide spectrum of depositional profiles and facies belts (Burchette et al. 1990; Burchette & Wright 1992; Badenas & Aurell 2001). The distinctive character of carbonate ramp systems shows the control of the physical, chemical and biological conditions that result from variations in palaeoclimate, tectonic regime, ecological changes, etc. Ramp systems are an important component of many carbonate successions, particularly during the early stages of the platform evolution (Read 1985). Ramps are widespread throughout the geological record and were particularly common during Mississippian and Jurassic times (Burchette et al. 1990; Burchette & Wright 1992; Badenas & Aurell 2001). Ramp systems contain significant hydrocarbon reserves. In particular, oolitic ramp systems form important reservoirs in various sedimentary basins of which the Upper Jurassic Smackover Formation of the US Gulf Coast is one of the best known (Heydari 2003). Other such examples are the Middle and Upper Jurassic reservoirs of the Persian Gulf: Marrat, Dhruma, Izhara and Araej Formations (Alsharhan & Kendall 1986; Alsharhan & Whittle 1995).

Although the factors controlling carbonate production and accumulation are relatively well understood, the analysis of ancient carbonate rocks is commonly challenged by the complexities of stratigraphic packaging and by incomplete or poorly exposed outcrops. A good knowledge of reservoir architecture (geometries and reservoir facies dimensions) is paramount for the construction of geological models and geostatistically-based reservoir simulations (Lomando 1998; van Buchem et al. 2002). Unfortunately, most of the models in literature (Aurell et al. 1995, 1998; Badenas & Aurell 2001; Burchette et al. 1990) are based on poorly exposed and/or discontinuous outcrops. This fragmentation compromises fine-scale observations where the precise facies geometries within high-frequency cycles need to be understood. This study presents the relationships between stratal and lithofacies anatomy in an almost completely exposed ooiddominated ramp system in the Lower to Middle Jurassic. The outcrops are located at the southern flank of the High Atlas in Morocco and have seismic scale dimensions (35 km long and about

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 65– 88. DOI: 10.1144/SP329.4 0305-8719/10/$15.00 # The Geological Society of London 2010.

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1000 m thick). A reliable biostratigraphic framework based on ammonites and brachiopods provides excellent time stratigraphic control on the scale of high frequency cycle (small scale) stacking patterns. This allows the analysis of the spatial distribution of oolitic facies and peloidal facies within high-frequency cycles as well as the evolution of depositional environments at different scales of time and space. The understanding of the controls on production and redistribution of ooids and pellets in oolitic ramp settings will help in predicting facies architecture in analogue reservoir rocks in the subsurface.

Geological setting The study area is located on the southern flank of the Central High Atlas mountain belt, approximately 30 km north of Goulmima (Fig. 1a), on both sides of the Gheris Valley (Fig. 1b). Uplifted during Cenozoic times, this mountain area evolved from an intracontinental basin called the High Atlas Trough (Ellouz et al. 2003; Arboleya et al. 2004; El Harfi et al. 2006a, b). During early Mesozoic times, this trough was a corridor open to the Tethyan Ocean that developed at the boundary between the Saharan Craton and the Moroccan Mesetas (Fig. 1b).

Fig. 1. (a) Map of the main structural domains of Morocco (modified after Pique & Michard 1989). The Rif is a mountainous arc that was thrust southward and dominates northern Morocco. The western and eastern Mesetas constitute individual microplates from Atlantic rifting. The Atlas thrust-belts extend 2000 km from the Atlantic Ocean to the Mediterranean in Algeria. The Anti-Atlas is the southernmost mountain belt of Morocco and is oriented ENE– WSW. (b) Tectonic and palaeogeographic diagram of the High Atlas basin during the Toarcian and location of the Amellago transect (compiled and modified from Laville 1985 and Milhi et al. 2002).

JURASSIC OOLITIC RAMP IN MOROCCO

According to numerous studies, the basin evolution can be summarized into two main tectonically-induced sedimentary phases, both linked to the Western Tethys and Central Atlantic rifting-drifting processes that occurred during the same period (Mattauer et al. 1977; Laville 1981; Ait Brahim et al. 2002). The first phase took place during north–south to NW–SE extension in the late Triassic to the late Lias when a true rift basin, created by NE–SW syn-sedimentary extensional faults, appeared at the northern boundary of the Saharan Craton. During the Triassic, welldeveloped half grabens were filled with tholeitic basalts, fluvial sandstones, and continental red beds and evaporites (Michard 1976; Pique & Michard 1989). During the early and middle Lias, the rifting phase continued with a rapid increase in accommodation space. Block tilting caused by high strain normal faults (Sarih et al. 2007) led to a major marine incursion from the Tethys Ocean and to a well-developed hemipelagic depocentre bordered by carbonate platforms (Wilmsen & Neuweiler 2008). The carbonate platforms are characterized by bioconstructed shelf breaks and steeply-inclined slopes along the borders of the basin or on the tilted blocks. At the end of the rifting phase, during the lower and the middle Toarcian, a eustatic rise of sea level caused a major drowning of all the regional carbonate platforms (Pique 1994; Elmi et al. 1999; Wilmsen & Neuweiler 2008). The second tectonosedimentary phase, that is, the post-rift evolution of the basin, began during the late Toarcian (upper Lias deposits), when the sinistral movement of Africa relative to Eurasia induced a transtensional regime. This led to the development of a mosaic of rhomb-shaped depocentres bounded by syn-sedimentary ridges (Laville 1988; Brede et al. 1992; Laville et al. 2004). From the late Toarcian to the late Bajocian, most of these depocentres were filled by hemipelagic marls, whereas carbonate platforms nucleated on the margins of the rhomb-shaped basins (e.g. the Amellago-Agoudim ramp system described here). In the late Bathonian or Callovian, the structural regime became transpressive and the High Atlas Trough was emergent. The Amellago transect is a series of cliffs and steep hillsides that form a quasi-continuous section across the Lias –Dogger platform-basin system (Fig. 2). The transect is oriented SSW–NNE and cross cuts the southern margin of a rhomb-shaped basin that developed between two syn-sedimentary faults, the Foum Zabel fault to the south and the Tagountsa fault to the north (Fig. 2). This seismic scale outcrop allows direct tracking of the stratal geometry and the facies migration that occurred along the palaeoslope.

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Two formations are well exposed in this transect (Figs 2 & 3). The Agoudim Formation (lower Toarcian to lower Bajocian) includes mainly blue and grey marls deposited in offshore environments (Hadri 1993). It forms the substratum, the cover, and the distal parts of the studied Amellago Formation (Fig. 3). The Amellago Formation includes all of the neritic carbonates that were deposited in the area from the late Toarcian to the early Bajocian. The Amellago Formation represents an oolitic carbonate ramp system that prograded northward into the subsiding basin (Figs 3 & 4) (Poisson et al. 1998; Durlet et al. 2001).

Methods Nearly 50 sedimentological sections (6400 m of cumulative thickness) were described for facies analysis and to collect biostratigraphic information. More than 350 thin sections were analysed petrographically and, along with the field observations, formed the basis for eight facies associations with a total of 22 facies types (Table 1; Figs 5, 6 & 7). The ‘depositional region’ terminology was modified from Tucker & Wright (1990), who defined an inner ramp and outer ramp, respectively located landward and seaward of the ooid shoal, which represents the middle ramp setting. Classically, the hemipelagic basin corresponds to the lower offshore zone deposits (Table 1). This terminology is well suited to the broad-scale progressive lithological changes that we observed between different environments along low-angle chronostratigraphic surfaces of the 2D strata framework. Measured sections were correlated by tracking high frequency sequences from inner ramp to basinal environments. Observations were recorded on photo mosaics including photos taken from the air in a microlight aircraft. Such photos have a resolution of about 10– 20 cm and were generally orientated perpendicular to the actual (dip direction) cliff face, instead of oblique from the foot of the cliff. Three high resolution cycles were selected for a detailed study because of their continuity and their accessibility in the field. The three cycles, each 20 to 30 m thick, have been traced in the field from the proximal to the most distal areas. Such correlations form the basis for the structural framework, the spatial and temporal facies relationships, the regional stratigraphic framework and the different facies models.

Facies description The facies are grouped in five main depositional environments comprising different sets of facies associations.

68 A. PIERRE ET AL.

Fig. 2. 3D view of Amellago transect showing the sedimentological series ages, the main sections and main geographic features (from Worldwind).

Fig. 3. Photo-montage of the Amellago transect with the formations and members interpreted in terms of major depositional environments. Vertical scale (3.5) is exaggerated because of the very flat ramp geometries.

Fig. 4. Stratigraphic correlation at the third-order scale from upper Toarcian to lower Bajocian. Am2 to Am5: aggrading-prograding Oolitic Ramp Sequences. The sequence Am2 is the first occurrence of a shallow carbonate platform in the area since the Pliensbachian. At the Djebel Ich Taskiouine measured section, sequence Am2 reaches 70 m of thickness. The oobioclastic facies (FA5) belt attains 8 –10 km of proximal-distal extension. At a smaller time scale, two high frequency cycles can be easily differentiated within Am2, both on pictures and on sedimentological sections. Sequences Am3 and Am4 show multiple toplap geometries, observed on the microlight pictures (see Fig. 8). Cumulative thickness of these sequences is 30 m on Djebel Ich Taskiouine but reaches more than 250 m northward, near Ksar Agoudim. Such thickness variation indicates that during this period, the northern part of transect was clearly subsiding more than the southern part. Elmi (1990) and Sadki (1992) already reported margin uplifts and basin subsidence. Oobioclastic facies (FA 4) are volumetrically dominant in sequences Am3 and Am4, which both show the farthest progradation. The resulting oobioclastic facies belt reaches a width of 20 km at the maximum of progradation. Sequence Am5 is a complete transgressive-regressive cycle. The prograding phase is accompanied by the development of oolitic grainstone facies (FA 4), which forms low angle clinoforms during this third-order highstand. The average extent of oolitic facies belts (shallow ramp facies) reaches 15 km during prograding phases and 8 km during retreating phases. Sequences Am6, Am7 and Am8: aggrading-retrograding muddy ramp sequences end the deposition of Amellago Formation during the Bajocian.

Table 1. Facies classification. The last last three columns display the main characteristics of the 22 lithofacies in terms of sedimentary structures and main Dunham (1962) textures Environment

Facies association and bathymetry

Facies (‘litho-facies’) Code

Inner ramp

FA1 Supratidal to intertidal þ2 to 22m

F1a

Tidal flats marls

Laminated

F1b F1c F1d

Mudflat M Grainy tidal flats W-P Restricted marine dolomicrite Lagoon marls Lagoonal micritic limestone W-F Lagoonal grain-dominated limestone P-G Storm wash over and tidal inlet P-G Coarsening-up M to G with mudclasts Marls Tide influenced shoal G-P

Laminated, mudcracks, microbial crusts Fenestral, mud cracks Laminated with sulphate pseudomorphs

Micrite, illite, chlorite, quartz and feldspar Micrite, benthic foraminifera Pelloids Dolomicrite, sulphates

Laminated, low energy Massive, low energy

Gastropods Large oncolites

Massive to laminated, moderate energy

Pelloids, aggregates

2D megaripples and cross beds, high energy Massive to laminated, burrows

Ooids

F2a F2b F2c F2d

Middle ramp

Outer ramp

FA3 ‘Open’ muddy lagoon 0 to 210m FA4 Oolitic belt

F3a F3b F4a

0 to 210m

F4b F4c

FA5 Upper offshore type 1 210 to 235m

F5a F5b F5c F5d

Structure and energy

Main content

Pellets, mudclasts, bivalve and coral clasts Laminated Brachiopods Tidal structures (e.g. sigmoids, climbing Pellets, ooids, bioclasts ripples, channels, bi-directional cross-bedding) Wave influenced shoal G 2D and 3D megaripples Ooids, bioclasts Coarse-grained shoal Massive and 2D megaripples Corals and shells fragments R-G Toe of shoal fine 2D ripples to massive, bioturbated Fine pellets, ooids, bioclasts grainstone Proximal tempestites Erosive base, scours, HCS Pellets, ooids, bioclasts R-G-P-W-M Nodular W-P Wavy, massive to laminated Pellets, ooids, bioclasts Patchy coral Massive Corals, bivalves, brachiopods, bioconstruction B echinoderms

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FA2 Lagoon 0 to 23m

Name

(Continued) 69

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Table 1. Continued Environment

Facies (‘litho-facies’) Code

FA6 Upper offshore type 2 220 to 250m

F6a

FA7 Hemipelagic basin 235m to 2150 m ? Ooid-free FA8 210m to muddy ramp 235m

F6b

Name

Heterolitic wavy beds W-P Distal tempestites R-G-P-W-M

Structure and energy

Nodular, Thalassinoides

F6c

Biostrome B

Erosive base, HCS normally graded beds (rudstone to mudstone with marls) Stratified

F7

Lower offshore to basin marls

Planar laminar, some slumps, low energy

F8

Glauconitic peloidal sand Heavily bioturbated, some scours and P HCS, generally low energy

HCS, hummocky cross-stratification; M, mudstone; W, wackestone; P, packstone; G, grainstone; R, rudstone; F, floatstone; B, boundstone.

Main content

Large bivalves (trichites, pholadomia and tridacnea) Pellets, ooids, bioclasts Corals, bivalves, brachiopods, echinoderms Rare ammonites, belemnites, zoophycos Foraminifera, pellets, micritic intraclasts, brachiopods, bioclasts, detritic quartz, wood debris

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Basin

Facies association and bathymetry

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Fig. 5. Inner ramp facies belt pictures. (a), Facies 1b, laminated mudstone with rip-up clasts; (b), Facies 1d, anhedral to subhedral dolomite anhedral, inclusion rich; (c), Facies 1c, fenestral texture, outcrop; (d), Facies 1c, fenestral texture, thin section; (e), Facies 1c, peloidal wackestone bed with mudcracks; (f), Facies 1a, coal debris (10– 20 cm) within marls; (g), Facies 2b, floatstone with oncoids.

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Fig. 6. Middle ramp facies belt pictures. (a), Facies 4a, tide-influenced shoal (note the size and the cyclicity of bundles; Jacob’s staff ¼ 1.5 m); (b), Facies 4a, tide-influenced shoal with composite stratification, decantation lamination and subordinate ripples; (c), Facies 4b, tabular cross-bedded oobioclastic grainstone; (d), Facies 4b, trough cross-bedded oobioclastic grainstone; (e), Facies 4b, grainstone detail with radial-fibrous ooids; (f), Facies 4c, grainstone with ooids, peloids and bioclasts; (g), Facies 4a, grainstone with superficial ooids, bivalve fragments and serpulids.

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Fig. 7. Outer ramp facies belt and basin pictures. (a), Facies 8, very fine packstone with peloids, echinoid fragments, dolomite, glauconite and quartz; (b), Facies 5c, wackestone with various bioclasts; (c), Facies 6c, bioconstruction at the offshore-basin transition; (d), Facies 6a, trichites and terebratula in life position; (e), Facies 6a and 7 alternations; (f), Facies 7, Hildoceras bifrons (ammonites); (g), Facies 7, hemipelagic marls.

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Inner ramp The inner ramp is located immediately landward of the ooid shoal belt and was a low-energy protected setting. The facies associations are dominated by micritic limestone, dolomicrite and marls which generally indicate low- to moderate-energy subtidal, intertidal and supratidal settings (Table 1 and Fig. 5). Facies Association 1: supratidal to intertidal environment. This facies association occurs in four intervals and is characterized by decimetre scale, tabular shaped beds with abundant ostracods, benthic foraminifera and gastropods. † Facies 1a: tidal flat marls: composed of red to green marls containing illite, chlorite, quartz, and feldspar. † Facies 1b: mudflat: consists of laminated mudstone with generally well-preserved mudcracks near the top of the intervals and locally contains rip-up clasts. † Facies 1c: grainy tidal flat: is dominated by peloidal grainstone to wackestone with peloids 0.25 and 4.0 mm in diameter. Fenestral fabrics and polygonal mud cracks are also common. They are indicative of sub-aerial exposure. † Facies 1d: restricted marine: is dominated by Fe-poor and inclusion-rich dolomite. The dolomite usually forms small euhedral crystals but large inclusion-free overgrowths may also occur. The paragenetic relationships indicate that dolomite precipitated prior to all other diagenetic phases. Small calcitic nodules that are 1 cm to 4 cm in diameter may occur within dolomite beds. Their occurrence between laminated dolomitic crusts, their shape, and their association with collapse breccias are indicative of ancient evaporite nodules that have been calcified. For all these reasons, the facies 1d dolomite is thought to have been deposited in a restricted marine environment (intertidal to supratidal), with intermittent phases of evaporation. Facies association 2: lagoon environment. The subtidal lagoon facies association occurs in three intervals of decimetre to metre scale and tabular bedding. † Facies 2a: lagoonal marls: is composed of marls containing very rare fauna (gastropods). † Facies 2b: lagoonal micritic limestone: consists of oncoidal wackestone to floatstone containing large micritic spheres with irregular cortices interpreted as oncoids (8 mm average diameter up to 2 cm). Vertical and horizontal burrow systems interpreted as Thallasinoides, probably of crustacean origin, are very common. † Facies 2c: lagoonal grain-dominated limestone: is composed of peloidal packstone to grainstone

with abundant well-sorted and well-rounded peloids and rare ooids. There are two types of peloids: faecal pellets and mud lithoclasts and their origin is often hard to determine. In some cases, the uniformity of their sizes and shapes is a clue that they are of faecal origin. Sometimes, peloids form aggregates with soft (deformed) intergranular contacts. The preservation and the burial of pellets are possible in this low energy environment. Also common are gastropods and vertical and horizontal burrows (Thallasinoides). † Facies 2d: storm washover lobes and tidal inlet grainstone: is characterized by coated-grain grainstones with a small proportion of bioclasts. It occurs in up to 1 m thick, 2D megaripple crossbeds that migrate landward (southward) and pinch out within facies 2a, 2b and 2c. The oobioclastic mega-ripples also form lenses detached from the oolitic shoal. The grainy beds are attributed to high-energy events and depositional sub-environments of storm washover lobes and tidal inlets that are located in an otherwise protected shallow marine environment behind the oolitic shoal. Facies association 3: open-marine muddy ramp environment. The subtidal, open-marine, muddy ramp association occurs in three decimetre to metre thick, tabular bedded intervals. † Facies 3a: mudstone to mudclast grainstone in graded beds: consists of beds with a mudstone base transitioning upward into wackestone with mudclasts followed successively by coarse, mudclast packstone and then mudclast grainstone. The bed tops are commonly bioturbated, bored, and iron stained. The mudclasts in the wackestone are somewhat deformed possibly by compaction during very early stages of burial. In addition to mudclasts, rare shell fragments of brachiopods and bivalves and even more rare coral fragments are observed. † Facies 3b: open-marine, muddy ramp marls: consists of marl deposits commonly associated with well-preserved brachiopods, which are indicative of low energy environments. The higher abundance of carbonate mud compared with the lagoon and the absence of a barrier or any oolitic shoal justifies the interpretation as an open-marine environment.

Middle ramp The shallow ramp area is dominated by oobioclastic grainstone (Facies association 4, Table 1; Fig. 6). When homogeneous shallow ramp grainstone facies of several high-resolution cycles are stacked, they constitute the most impressive cliffs of the Amellago transect (up to 200 m thick).

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Facies association 4: oolitic belt environment. Facies association 4: Oolitic belt: consists of three decimetre to metre scale, alternating, mainly tabular but also lobe-shaped facies bodies. Based on sedimentary structures, three facies types are distinguished. † Facies 4a: tidal influenced shoal: is characterized by bi-directional cross-beds, alternations of fine and coarse bedsets showing a cyclic organization, and sigmoidal cross stratification typical of tidal environments. Facies 4a is interpreted as tidal oolitic sand bars in a tidal inlet. † Facies 4b: wave influenced shoal: is mainly composed of grainstone and exhibits 2D and 3D mega-ripple cross beds and low angle cross beds that appear to prograde seaward. Reactivation surfaces are present. The deposits are organized along low angle (up to 3–48) clinoforms, which also show a seaward migration. Facies 4b is interpreted as shoreface sands. † Facies 4c: coarse shoal facies: consists of massive deposits of coarser bioclastic debris (molluscs, echinoderms, and corals). In general, Facies association 4 is dominated by oobioclastic grainstones with subordinate packstones. Grainstones are composed of ooids and bioclasts with 70% of non-skeletal grains. Three types of ooids can be differentiated based on thin-section analysis: (1) radial-fibrous ooids with several regular cortical layers; (2) superficial ooids with a thin cortical layer; and (3) small micritic ooids (Fig. 6). Nuclei of radial-fibrous ooids and superficial ooids are mostly peloids, but can also be bioclasts, aggregate grains and lithoclasts. There are also composite ooids in which the nucleus consists of one or several ooids held together by a micritic matrix. Ooid cortices have a radial crystal structure interrupted by thin, dark micritic layers. Some peloids that form nuclei are dolomitized and stained red by iron oxides, facilitating identification with a hand lens. Bioclasts (echinoderms, bivalves, brachiopods, gastropods, bryozoans and foraminifera), lithoclasts and aggregate grains are quite common within the grainstones. Sorting and concentration of the different components depends on the position within a high-resolution cycle. The described sedimentary structures and grain-rich nature of this facies association indicate deposition on a high-energy, ooid shoal under the influence of fair weather waves and local tides.

Outer ramp Facies association 5: proximal upper offshore environment. Facies association 5 contains four different facies types with variable stratigraphic organization.

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† Facies 5a: toe-of-shoal fine grainstone: is dominated by very fine to fine, well-sorted, and wellrounded peloidal grainstone to packstone. The abundant peloids with an average diameter between 0.20 and 0.5 mm have a distinctive round shape. Their regular size, shape and internal concentric structure suggest a faecal origin. Other common allochems are ooids and bioclasts. Sedimentary structures such as ripples are often destroyed by intense bioturbation, which suggests that sedimentation rates were probably reduced and the environment was at least intermittently subject to lowenergy conditions. † Facies 5b: proximal tempestites: consist of fining upward, decimetre-scale, lensoidal beds (rudstone to mudstone) that are 10 to 70 cm thick. The beds are commonly amalgamated with erosive bases and usually show hummocky cross stratification (HCS) in the proximal part. These beds are interpreted as tempestites in the sense of Aigner (1985) and deposited between fair weather wave base (FWWB) and storm wave base (SWB). Graded beds interfinger with bioturbated limestone and marls just below the SWB in the distal part. † Facies 5c: nodular wackestone packstone: is composed of peloids and bioclasts (brachiopods, bivalves and echinoderms fragments). In distal areas, coquinas are found scattered or concentrated in pockets and are associated with erosive features with normally graded bioclasts and/or HCS. These deposits are interpreted as tempestite relicts that were reworked by bioturbation. The depositional environment is interpreted as a weakly agitated area from distal shoreface to proximal offshore. † Facies 5d: coral bioconstructions: is composed of framestone to bafflestone and occurs in patches of about 5– 10 m width and 1–2.5 m height. Hermatypic corals with mainly branching and domal forms and sponges form the framework of the buildups. Corals are rarely reworked indicating deposition below the FWWB. Buildups formed within tempestite deposits and are interpreted to occur in the photic zone in the proximal part of the offshore environment. Facies association 6: distal upper offshore environment. The offshore facies association type 6 (distal) consists of 3 different facies types with various natures and various origins. † Facies 6a: heterolithic wavy beds: is composed of alternations of tabular 2–5 cm thick beds of marls and 2–10 cm thick beds of mudstone to wackestone. Thalassinoides traces attest that the nodularity probably results from the combination of heterogeneous bioturbation and

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compaction. The presence of large bivalves (trichites, pholadomia and tridacnea) usually in life position attests to an open-marine, distal depositional environment. † Facies 6b: distal tempestites: consists of normally graded beds (rudstone to mudstone) of 10–70 cm thickness. This facies is similar to Facies 5b with the exception that the tempestites beds are interbedded with thin layers of marl. The lower limit of the deep ramp is marked by the disappearance of the graded tempestite beds. † Facies 6c: stratified bioconstructions: is characterized by columnar-shaped facies bodies with average dimensions of 3 m width by 6 m thickness. In exceptional cases, biostromes of 100 m width and 20 m thickness were observed. The bioconstructions are composed of hermatypic corals of various forms, sponges, and large bivalves (tridacnae) that are slightly broken up. These carbonate masses are formed by successive layers slightly thickened (few decimetres) that differ from surrounding Facies 6a rocks. Relationships with surrounding layers show that: (1) their palaeo-elevation was low, on the order of several decimetres above the seafloor; and (2) they grew in the distal part of the upper offshore zone and formed in water depths of a few tens of metres.

fauna are mainly composed of ammonites, small bivalves and small gastropods.

Ooid-free muddy ramp Facies association 8: Inner to mid ramp environment. Facies association 8 is represented by one facies. It forms homogeneous argillaceous microwackestone to micro-packstone prisms that are several kilometres in extent and with a thickness of several decimetres to a metre. † Facies 8: glauconitic peloidal packstone: is composed dominantly of small peloids and accompanied by common glauconitic and quartz grains, wood debris, and microbored bioclasts including bryozoans, foraminifera, brachiopods, gastropods, and bivalves. Radialfibrous ooids and micritic ooids are never present, whereas superficial ooids and ooid fragments are very rare. This facies is often dolomitized, especially in the proximal parts of the prisms. The beds are structureless, which is interpreted as a sign of intense bioturbation. The weathering gives a light beige colour to this ooid-free enigmatic facies (see part 5).

Evolution of the ramp system at the large scale

Hemipelagic basin

Stratigraphic architecture

Facies association 7: Lower offshore to basin environment. Facies association 7 is represented by one homogeneous facies. † Facies 7: Lower offshore to basinal blue-grey marls: is composed of blue-grey marls with an average carbonate content of 20 –30%. Others components are clays (illite, chlorite) and fine terrigenous grains (quartz and feldspars). The

The migration of the facies belts seaward (progradation) and landward (retrogradation) are easily identifiable in the Amellago transect. On the basis of their internal progradational, aggradational and retrogradational geometries (Fig. 3), the Agoudim and Amellago Formations can be subdivided into a hierarchy of three orders of cycles, large, medium and small (Durlet et al. 2001) (Table 2).

Table 2. Biostratigraphic zones for stratigraphic surfaces with numerical ages from Gradstein et al. (2004) Surfaces and medium scale cycles

Ammonites biozones from Groupe Franc¸ais d’etude du Jurassique (1999)

DS9

Discites/Laeviuscula/ Propinquans

DS8–Am8 DS7–Am7 DS6–Am6 DS5–Am5 DS4–Am4 DS3–Am3

Concavum/Discites Concavum Concavum Murchisonae/Bradfordensis Opalinum (Bifidatum) Opalinum

DS2–Am2

Meneghini/Aalensis

Calculated numerical ages for stages boundaries from Gradstein et al. (2004)

Stages

BAJOCIAN 171.6

175.6

base BAJOCIAN AALENIAN AALENIAN AALENIAN AALENIAN AALENIAN AALENIAN base AALENIAN TOARCIAN

JURASSIC OOLITIC RAMP IN MOROCCO

Large-scale cycle. There is only one large-scale cycle (Fig. 4). The older (first) hemicycle starts within the thick deposition of several hundred metres of hemipelagic marls (Agoudim Formation) and finishes at the maximum basinward migration of the Amellago formation dominated by oolitic facies (FA4). The following hemicycle (second) is characterized by the landward migration of the Amellago Formation without any FA4 facies. Deposits from the retrograding and aggrading mud-dominated depositional systems lie on top of the oobioclastic ramp deposits. Medium-scale cycles. At a higher frequency (medium-scale cycles have a thickness 50– 100 m), five cycles of migration of facies belts seaward and landward along a dip section have been identified in the first large hemicycle (Am1 to Am5; Fig. 4). The surfaces at the bases of these cycles are noted as DS1 to DS5. Three medium-scale cycles are identified in the second large hemicycle (Am6 to Am8). The surfaces at the bases of these cycles are noted as DS6 to DS8. The top of Am8 is called DS9. Cycles Am1 to Am5 each contain oolitic facies (FA4). Subsequently, an abrupt change in facies occurs near the Aalenian –Bajocian boundary, at the base of sequence Am6 (DS6). Above the DS6 surface, the successive environmental belts from the proximal to the distal ramp are represented by facies associations FA1, FA3, FA8, FA5 and FA7. The sedimentary system was a ramp with ubiquitous muddy textures but without ooids along the proximal distal transect. Small-scale cycles. The five medium-scale cycles of the oldest large-scale hemicycle can be subdivided into 26 small-scale cycles (10 –30 m thick; Fig. 4). Each small-scale cycle is bounded at its base by a discontinuity surface showing similar lateral sedimentological and diagenetic variations which will be described in ‘Spatial and temporal distribution of facies within high frequency cycles’ section. In the youngest large-scale hemicycle, the medium-scale cycle 6 contains three small-scale cycles. The small cycles within medium-scale cycles 7 and 8 are difficult to observe and are not discussed in this paper.

Biostratigraphy and dating The geological period studied here, from Toarcian and to Bajocian, is well dated by an abundant, diverse ammonite and brachiopod fauna (Durlet et al. 2001; Almeras et al. 2006; Pierre 2006; Bourillot et al. 2008). Correlations between measured sections as well as the timing of significant stratal

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surfaces are constrained by biochronological charts (Groupe Franc¸ais d’Etude du Jurassique 1997; Almeras et al. 2006). The numerical ages of these timelines are calibrated to the timescale of Gradstein et al. (2004). The ages of the medium-scale cycles are, respectively: Menegheni/Aalensis zones (late Toarcian) for Am1, Opalinum zone (Aalenian) for Am2 and Am3; Murchisonae zone (Aalenian) for Am4, Concavum zone for Am5 and Discites/Laeviuscula/Propinquans zones for Am6, Concavum zone for DS7, Concavum/Discites for DS8, and Discites/Laeviuscula/Propinquans for DS9 (Durlet et al. 2001; Pierre 2006).

Sequence stratigraphic and cyclostratigraphic interpretation The sequence or cyclostratigraphic interpretation of the Amellago/Agoudim ramp system is based on the organization of the migration of facies belts landward or seaward along a dip section. The cycle or sequence definition is based on the notion of geological time. They are the result of tectonic, tectono-eustatic or glacio-eustatic mechanisms (Vail et al. 1991). The evolution of ramp system follows this hierarchy of cycles (large, medium and small). The large-scale hemicycles are several hundred metres in thickness and record deposition over several millions of years (Ma) (Table 2; Fig. 4). They are considered second-order sequences (Vail et al. 1991). The first hemicycle is defined by the aggradation of Lower Toarcian hemipelagic marls (Agoudim Formation) and the northward progradation of the Amellago Formation dominated by oobioclastic ramps systems from the Toarcian – Aalenian transition, during the entire Aalenian and approximately to the Aalenian –Bajocian boundary (Sequences Am1 to Am5). The second hemicycle is defined by the southward aggradation – retrogradation of the Amellago Formation southward from the Aalenian –Bajocian transition to the middle Bajocian (Sequences Am6 to Am9). The medium-scale cycles have an average thickness of 50–100 m (Fig. 4) and represent an average duration of 0.5–3 Ma. (Table 2). The updip and downdip shifts in facies zones (landward and seaward respectively) suggest the alternation of transgressive-regressive cycles (Fig. 4). They are considered third-order sequences (sensu Vail et al. 1991). A high frequency cycle is defined as a composite transgressive-regressive interval bounded by two successive time surfaces of the same origin (Homewood et al. 1992). Thus, the small-scale cycles are considered high-frequency cycles (Fig. 4).

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Spatial and temporal distribution of facies within high frequency cycles Description of cycle architecture This adopted methodology allowed the documentation of the lateral distribution of facies and discontinuities and their time relationships within three high-frequency cycles (Am2a, Am4a and Am4b) (Figs 8 & 9). The evolution of the depositional model along timelines within each high frequency cycle is described below. Basal discontinuity surface. The basal discontinuity surface has been physically correlated and tracked. In the inner ramp environments, the surface is generally represented by a hardground that usually overlies and truncates the intertidalsupratidal dolomitic facies of the preceding high frequency cycle. The hardground forms a surface with irregular relief, local encrusting fauna and bivalve borings. Early cementation is indicated by micrite cementation and by thin isopachous calcite cements around allochems. This cement is inclusion-rich and shows a cloudy luminescence which suggests a high magnesium calcite (HMC) precursor (Moore 2001). It was probably precipitated from seawater in a phreatic environment during the sedimentary gap associated with the unconformity. Neither vadose and/or meteoric cements nor karstic or pedogenetic structures have been observed below or above this basal discontinuity surface. In the mid-ramp, the surface is a planar hardground with numerous borings attributed to bivalves, echinoids and annelids. Iron hydroxides and encrusting fauna (oysters, bryozoans, and serpulids) are also present. Its lateral extent, from proximal to distal areas, reached more than 10 km with an average slope of 0.1–0.028. Early lithification of the underlying oolitic grainstone is shown by the presence of isopachous fibrous cements, which are truncated by borings and by the planar surface itself. These calcitic cements are inclusion-rich and exhibit a cloudy luminescence indicative of HMC precursor cement. As shown in other studies (Dravis 1979; Loreau & Durlet 1999; Moore 2001), these characteristics suggest marine cementation and abrasion in an agitated shoreface environment. In this mid-ramp setting, the surface is covered by glauconitic calcisiltite FA8 or by hemipelagic marls FA9 in sharp contrast with the underlying oolitic grainstone. Basinward, the planar hardground shifts gradually to a bioturbated and non-lithified surface, and dissipates into an inconspicuous, conformable bedding surface within the hemipelagic marls of the basin. Ooid-free interval. In this interval, the physical correlation of stratigraphic beds shows a depositional

profile that is planar and slightly inclined (0.1– 0.028) without any slope break (Fig. 8). The entire ramp profile is dominated by muddy, peloidal and dolomitized lithofacies, and ooids are completely absent. The successive environmental belts from the proximal to the distal ramp are represented by facies associations FA1, FA3, FA8, FA5 and FA7. The argillaceous peloidal very fine sands (Facies 8 packstone) form a prism, which pinches out seaward (Fig. 8). The planar upper surface of this prim is also an erosive surface. The flooding during the transgression and the associated hydrodynamic changes (razor effect, sensu Eichenseer & Leduc 1996) may explain this. Brachiopods are relatively common in these transgressive ooid-free facies, even in the inner part of the muddy ramp which confirms their preference for such open and moderately agitated environments (Almeras & Faure 1990; Almeras et al. 1994). Aggrading oolitic interval. The physical correlation of stratigraphic beds shows progressively greater inclination of timelines upward (Fig. 8). The large scale aggrading and slightly prograding geometries of the depositional oolitic system during this stage are typical. During this interval, the lateral facies succession along these time lines from a proximal to a distal environment is FA1, FA2, FA4, FA5, FA6 and FA7 (Table 1). The transitions between all the facies zones are gradational and make discrete interpretation difficult. The ramp slope and therefore the proximal– distal dimension of the middle ramp facies belts evolve in this interval from 0.028 degrees and several kilometres to a maximum of 38 and few tens of metres. Near the base of this unit and especially in the inner ramp domain, ravine surfaces are often observed. These irregular erosion surfaces mark a granulometric jump from fine facies below to coarse facies above the surfaces. They are interpreted as submarine erosion surfaces created by episodic high energy events probably created by hydrodynamic waves or tides. In the outer ramp domain, tempestites seem to be more important during this stage of inclined profile than at any other stage of ramp evolution. The outer ramp during this interval is also notable for the development of bioconstructional deposits (Facies 6b). The carbonate production (e.g. ooids) is generally localized in tidal oolitic systems and in permanently wave-agitated areas of the mid-ramp. Meanwhile, the transition zone between lagoon and oolitic shoals is a place of accumulation of tidal and storm wash-over deposits (Facies 2d). Differential rates of sediment accumulation along the slope profile lead to progressively steepening depositional profiles during this time, shown on photos by stratal surfaces that reach 3–48 at the slope break.

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Fig. 8. Correlation panel showing the lateral distribution of facies and discontinuities and their time relationships within high-frequency cycles 2a. The evolution of the depositional model along timelines is deduced from these correlations.

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Fig. 9. Progradation of oobioclastic clinoforms.

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Prograding oolitic interval. The physical correlation of beds shows that at the large scale, this interval is composed of prograding oobioclastic wedges. Such prograding wedges are mainly composed of low angle clinoforms with a maximum average slope of 38 (Fig. 9). The correlation of sedimentological sections as well as the analysis of photomosaics show that the facies succession along a time-line from proximal to distal environments is FA4, FA6 and FA7 (Figs 8 & 10). This succession includes ooid ramp facies; however, the restricted, proximal environments (FA1 and FA2) were either not deposited or not preserved. In addition, observed structures (cross beds, massive megaripples) in oolitic deposits reflect the most highly agitated hydrodynamic conditions within the high frequency cycle. Interpretation of the cycle architecture. The high frequency cycles are composed of laterally shifting (or ‘retreating’, aggrading, and prograding) facies tracts. Each of these tracts is associated with a particular sedimentary system. A conceptual model that illustrates the main facies elements, geometries, and early diagenetic properties of the 26 elementary high frequency cycles is presented in Figure 10. Transgression across a partially exposed platform. Based on the surface characteristics (early marine cementation, colonization by marine organisms) and the proximal facies observed below and the distal facies observed above, the basal discontinuity of the cycle is interpreted as the result of transgression across a partially exposed platform (Fig. 10). Start up of the carbonate factory, low energy ramp. Ooids are completely absent across the entire ramp during the early transgression. From the most proximal environment to the hemipelagic basin, mudstone, peloidal, oncoidal wackestone, peloidal and intraclastic packstone and grainstone, and marls dominate the muddy ramp system (Fig. 10). The proximal facies are very close to those of the oolitic system, except that there are no backshore oolitic sands. All across the muddy ramp system, the high energy facies are rare. They mainly include bioturbated grainstone and packstone with peloids, benthic foraminifera, micritic intraclasts, wood debris and glauconitic grains. The facies variation between proximal and distal muddy facies is transitional and subtle. The distal muddy ramp environment is one of intercalated marl-limestone interlayered with tempestites. The latter is close to the oolitic outer ramp but with more pelloids and no ooids. It appears that this ramp profile is very flat without any slope break. In contrast to the oolitic ramp system, this system is named ‘muddy ramp

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system’. All the environments reflect weakly agitated hydrodynamic conditions that were not favourable for ooid formation. Hydrodynamic changes induced during initial transgression (the wave razor) lead to multiple erosive surfaces like the planar upper surface of the pelletoidal very fine sand shoals (Facies 8). Higher accommodation, higher energy, aggrading ooid ramp. During the aggrading time interval, the ooid factory is restored and develops an oolitic ramp shoal (Fig. 10). The classic facies juxtaposition from inner ramp to the basin (Burchette et al. 1990; Burchette & Wright 1992; Badenas & Aurell 2001), the dimension of the facies belts and the depositional slope profile allow for the comparison with the present-day Persian Gulf Trucial Coast oolitic ramp systems (Loreau & Purser 1973; Loreau 1982). Reducing accommodation, infill and progradation of high energy ooid belts. During this phase of the cycle, ooid grains dominated, leading to a major progradation of the sedimentary system. One of the most striking observations is the extensive progradation driven by a surge in ooid production probably coinciding with and linked to a decrease in accommodation space inferred from beds thinning upward (Fig. 8). Low angle clinoforms and the partitioning of facies along a time line suggest that the belt of ooid production and/or deposition is narrow with a width less than 1 km. However, the lateral extension of the ooid facies bodies can reach more than 15 km (Fig. 3). Consequently, at the scale of the Amellago transect, these oolitic facies bodies are diachronous and the oolitic wedges are platform prisms which prograde and overlie the deep ramp deposits. End of cycle. The inner ramp may have been partially exposed. The cycle ends with a new surface. This surface is similar in terms of shape and sedimentological characteristics to the basal surface. Both surfaces must have had the same origin. They are the result of a transgression across a partially exposed platform. Above this surface, a new ooid-free muddy ramp system develops. This story is repeated 26 times in the studied interval.

Comparison with the Persian Gulf On the basis of modern oolitic systems from the Persian Gulf and British West Indies, some authors (Loreau & Purser 1973; Loreau 1982; Lloyd et al. 1987) suggest that ooid production and depositional belts are limited in width. Following this interpretation, the ancient, broad oolitic bodies could be the result of migration of the ooid factory or a

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Fig. 10. 3D diagrams showing evolution of depositional environments.

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Fig. 11. ‘Synthesis’ with characteristics of the two different ramp depositional systems during high frequency cycles. During transgression, low carbonate production rates with a low angle, a muddy system is established. Whereas during late transgression to late highstand, a phase of high production rates with an ooid dominated higher angle ramp system is established.

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redistribution of ooids from their production area. In the Holocene analogue of West Caicos, a beach and shoreface oolitic system can prograde over several kilometres in less than a million years and produce a sheet of oolitic sand averaging 5 m thick (Lloyd et al. 1987). The observations in recent environments and in the Amellago example presented here, suggest that continuous oolitic geobodies across great down-dip distances in other basins may be diachronous and the result of strong progradation and seaward migration of facies belts.

Discussion The Amellago ramp system is characterized by high frequency cyclic alternations of muddy early transgressive and ooid-dominated late transgressive and highstand intervals (Fig. 11). Although ooiddominated ramp systems are common (Burchette et al. 1990; Burchette & Wright 1992; Badenas & Aurell 2001), the presence of the muddy transgressive intervals has rarely been documented and may be linked to the specific factors controlling cyclicity at this ramp. For Strasser et al. (1999), climatic and eustatic cycles cannot be separated. The example of Kimmeridgian carbonate ramps of the Iberian Basin suggests that the relative sea-level changes are probably associated with climatic change and influenced the carbonate factory (Badenas et al. 2003, 2005). However, as Schlager (2005) mentioned, it is important to understand how these parameters act on a sedimentary system (e.g. sedimentary production change and morphology changes). The following sections discuss the relative influence of relative sea-level fluctuations and climatic changes on the high frequency cycles.

Eustacy-driven cyclicity Relative sea-level variations are known to have a significant impact on ooid production by directly influencing hydrodynamics and sea-water chemistry. Since the early seventies, hydrodynamic agitation has been well known as a key factor in ooid formation (Bathurst 1971; Loreau & Purser 1973; Lucas et al. 1976; Loreau 1982; Tucker & Wright 1990). It stimulates seawater degasification and contributes to evaporation and air-bubble formation, which are processes that increase supersaturation and CaCO3 precipitation (Girou 1970). Generally, ooid sands are abundant in such agitated environments while, in contrast, peloids are more common in calm environments. All facies zones observed in the muddy ramp system reflect weakly agitated hydrodynamic conditions. The increase in ooid fraction of facies is coincident with the evolution of more steeply

dipping depositional profiles. Schlager (2005) observed that ooids are more frequent during HST because tidal currents and wave action enhances their production. The influence of sea floor morphology on high-energy facies distribution was suggested by Purser (1983). During early transgression on the Amellago ramp, a very low angle depositional profile of about 0.018 was maintained and probably acted as a wide distal surf zone that would have dissipated all wave energy (Fig. 11). In addition, the low topographic relief and absence of any barrier such as ooid shoals is interpreted to have caused diminished tidal current strength, leaving the inner muddy ramp more exposed to the sea than the inner ooid ramp domain (Fig. 7c). Brachiopods are present in Facies 3 and indicate a more open marine condition (Almeras et al. 1994). As a result, salinity and carbonate saturation may have been lower due to a less evaporation than on a platform with a barrier. This, in turn, may have hindered ooid production. Low amplitude relative sea-level variations were sufficient to drown the Amellago ramp with facies belts shifting several kilometres. Eustatic sea-level variations can be explained by climate changes linked to astronomical Milankovitch cycles that influence insolation patterns on Earth (Vail et al. 1991). These changing insolation patterns controlled the sea level during the Pleistocene ‘icehouse’ by repeated melting and formation of polar and continental ice caps (Strasser et al. 1999). During the ‘greenhouse’ of Toarcian, Aalenian, and Bajocian times, presence of continental and polar ice caps is still under debate (Hesselbo & Jenkyns 1998; Cobianchi & Picotti 2001; Hallam 2001; Immenhauser 2005). The thermal extension/contraction of the upper layer of the oceans, the thermal change of volume in the deep oceanic currents, and the capture of water in aquifers and lakes (Gornitz et al. 1982; Menard et al. 1995; Cazenave 1999; Miller et al. 2005) are others possible factors that may have led to sea-level changes of high frequency low amplitudes during the Lias – Dogger transition. However, the question remains whether a rapid eustatic drowning of the ramp and the associated topographic change of the sea-floor could have triggered the total disappearance of ooids during transgression.

Climate-driven cyclicity Modern analogues in the Persian Gulf, West Caicos, and the Bahamas show that agitation is not the unique and only factor that controls ooid formation (Bathurst 1971; Loreau & Purser 1973; Lucas et al. 1976; Loreau 1982; Lloyd et al. 1987; Tucker & Wright 1990). In the Persian Gulf, the climate is hot and highly evaporative. Ooid

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formation here is significantly controlled by physical-chemical parameters such as temperature, salinity, pH and concentrations of Ca, Sr, and Mg that control the thermodynamics and kinetics of CaCO3 precipitation. In Amellago, a constant thermal subsidence inherited from the post-rift period creates accommodation space (Ellouz et al. 2003). If this space created by subsidence is not filled in by sediments, the ramp would drown. Consequently, the transgressive surface and the development of the muddy ramp system would result from a climaticallycontrolled carbonate and ooid productivity crisis. During the return of more favourable climatic conditions, ooid production could resume and fill the empty accommodation space until the next climatic crisis. Therefore, ooid production could also be predominantly controlled by climatic cycles, particularly arid versus humid cyclicity. In the Amellago case study, two observations support the hypothesis of a more humid climate during ooid-free intervals. (1) Locally, terrigenous particles in the muddy ramp deposits form thin beds of sandstones with a dominant micritic and locally microdolomite rich matrix, which have been observed westward of the Amellago ramp, near the deltaic systems of the Central High Atlas (Ettaki et al. 2000; Ettaki & Chellaı¨ 2005). Their occurrence in the Amellago area during ooid-free intervals could be the sedimentary signature of increased weathering of the Saharan Craton, probably during more humid climatic conditions. However, the presence of such terrigenous elements may also be the result of a change in the littoral drift trajectory when oolitic shoals and barriers were absent. Alternatively, clastic-rich intervals may represent lowstand or late highstand clastics that were reworked during the ensuing early transgression (e.g. Eichenseer & Leduc 1996; van Buchem et al. 2002). (2) Coaly debris, though rare is present in the muddy ramp deposits. This may indicate the development of forests or mangroves during relative humid periods (Tucker & Wright 1990). (3) Unfortunately, geochemistry (in particular d18O measurements on calcitic shells) cannot be used in the Amellago ramp system due to the burial history of this area; signatures are not preserved (Railsback & Hood 2001). As a result it is difficult to prove climatic versus eustatic control over the high frequency ramp system.

Conclusions The analysis of the carbonate system of Amellago leads to the recognition and characterization of geometries and the internal architecture of Jurassic

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oolitic ramp deposits. These results are afforded by the exceptional, seismic-scale dimensions of the continuous outcrop (37 km long and 1000 m high) along the ramp’s dip profile. Important insights include ramp slopes, palaeobathymetry (Table 1) and widths of facies belts estimated directly from the interpretation of a photo-mosaic profile. The detailed stratigraphic framework establishes that, at the large scale (several millions years and several hundred metres of thickness), sedimentary systems gradually shift from overall prograding oobioclastic ramps to overall retrograding ‘muddy ramps’ from the Toarcian–Aalenian transition through the entire Aalenian. At the small-scale (high frequency order cycle), correlation of serial sedimentological sections shows that the ramp profile evolved though time with the following: (1) The high frequency cycles are composed of laterally shifting (or retrograding, aggrading, and prograding) facies tracts. Each of these tracts is associated with a particular sedimentary system. (2) The retreating period is represented by ooidfree muddy systems and their aggrading and prograding intervals are represented by oolitic ramp systems. (3) The strong volumetric partitioning observed here suggests that continuous oolitic facies along great distances in other basins may be diachronous and the result of strong progradations and seaward migration of facies belts. (4) Depositional slopes and the widths of facies belts evolve over the course of a high frequency cycle (from 0.028 and several kilometres to a maximum of 38 and few tens of metres for the defined ‘middle ramp’). Cycles can be controlled directly by eustacy or more indirectly by variability in carbonate production, which is largely a function of climatic variations. It is believed that during the evolution of a cycle, ooid production progressively increases leading to a dramatic progradation of the sedimentary system. This unique outcrop is a reference for oolitic ramp models at several scales. This study is a project of ‘Equipe Stratigraphie quantitative et diagene`se’ UMR CNRS Bioge´osciences 5561, Universite´ de Bourgogne, France. Very special thanks to Y. Alme´ras, J.-L. Dommergues, S. Elmi, P. Neige, R. Bourillot and L. Rulleau for biostratigraphy. Warm and hearty thanks to C. Grelaud, J. Tranier, S. Ve´drine and M. Ousri for their help on the field. We are also glad to thank the Faculty of Sciences Semlalia (Marrakech) for the scientific and logistical help. During the final revision of the paper, we received helpful comments from J. Kenter, A. Saller, J. Hsieh, A. Reed and other Chevron colleagues. We also thank the reviewers for their very constructive comments.

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S ARIH , S., Q UIQUEREZ , A., G ARCIA , J. P., E L H ARIRI , K., A LLEMAND , P. & C HAFIKI , D. 2007. Se´dimentologie et quantification de la subsidence des se´ries liasiques dans le Haut Atlas Central marocain (coupe de Foum Zabel, Re´gion de Rich). Africa Geoscience Review, 14, 181–194. S CHLAGER , W. 2005. Carbonate Sedimentology and Sequence Stratigraphy. SEPM Concepts in Sedimentology and Paleontology, 8. S TRASSER , A., P ITTET , B., H ILLGARTNER , H. & P ASQUIER , J.-B. 1999. Depositional sequences in shallow carbonate-dominated sedimentary systems: concepts for a high-resolution analysis. Sedimentary Geology, 128, 201 –221. T UCKER , M. E. & W RIGHT , V. P. 1990. Carbonate Sedimentology. Blackwell Scientific Publications, Oxford. V AIL , P. R., A UDEMARD , F., B OWMAN , S. A., E ISNER , P. N. & P EREZ -C RUZ , C. 1991. The stratigraphic signatures of tectonics, eustasy and sedimentology: an overview. In: E INSELE , G., R ICKEN , W. & S EILACHER , A. (eds) Cycles and Events in Stratigraphy. Springer-Verlag, Berlin, 617 –659. VAN B UCHEM , F. S. P., P ITTET , B., H ILLGARTNER , H., G ROTSCH , J., A L M ANSOURI , A. I., D ROSTE , H. & O TERDOOM , W. H. 2002. High-resolution sequence stratigraphic architecture of Barremian/Aptian carbonate systems in northern Oman and the United Arab Emirates (Kharaib and Shu’aiba Formations). GeoArabia, 7, 461– 500. W ILMSEN , M. & N EUWEILER , F. 2008. Biosedimentology of the Early Jurassic post-extinction carbonate depositional system, central High Atlas rift basin, Morocco. Sedimentology, 55, 773–807.

Sedimentary evolution of an Upper Jurassic epeiric carbonate ramp, Iberian Basin, NE Spain ´ DENAS, J. IPAS & J. RAMAJO M. AURELL*, B. BA Dipartimento Ciencias de la Tierra, Universidad de Zaragoza, 50.009-Zaragoza, Spain *Corresponding author (e-mail: [email protected]) Abstract: A synthesis of the sedimentary evolution of the Upper Jurassic carbonate epeiric ramps that developed in the northern part of the Iberian Basin (NE Spain) is presented. The facies distributions reconstructed from the analysis of a 200 km-long transect, show a transition from shallow to relatively deep sedimentation sites. The studied carbonate ramps record major long-term transgression, from mid-Oxfordian to mid-Kimmeridgian followed by progressive basinwards coastal shift until the major regressive event around the mid-Berriasian. Subsidence was relatively homogeneous across the northern Iberian Basin during most of the studied interval. Major episodes of differential subsidence occurred around the Oxfordian– Kimmeridgian transition and onwards from the mid-Tithonian. The sedimentary evolution and particular facies types of the successive Iberian carbonate ramps is described, considering five depositional sequences that have a long-term transgressiveregressive evolution. The Oxfordian sequence shows a sharp transition from shallow to deep ramp areas: from mixed siliciclastic-carbonate (ooidal, skeletal, peloidal) facies to condensed (i.e. spongiolithic, peloidal, glauconitic) facies in the open platform domain. In the two Kimmeridgian sequences (Kim1 and Kim2), the transition between shallow and deep areas is more gradual and thickness distribution across the ramp is more homogeneous. Coral-microbial reefs and oolitic-peloidal-skeletal shoals characterized the shallow areas. Towards the offshore domain, these facies grade rapidly to a tempestite-dominated lithofacies and then into thick lime mudstone successions (i.e. rhythmic marls-mudstone alternations). Shallow oncolitic-peloidal and skeletal facies covered wide areas of the carbonate ramp during the early Tithonian (Ti1 sequence) and graded basinwards to thick successions of well-bedded micrites. The middle Tithonian to lower Berriasian platform (Ti2 sequence) is only partly exposed and formed during a stage of more heterogeneous subsidence. It is characterized by a thick succession with metrescale shallowing-upward sequences with local development of peritidal, algal-laminated caps. The factors that controlled the sedimentary evolution and major facies changes across the successive epeiric carbonate ramps are discussed by comparison to other Upper Jurassic platforms developed in the western Tethyan realm.

Epicontinental seas spread over wide cratonic areas during greenhouse periods, when sea-level was in a much higher position than today. This resulted in the development of epeiric carbonate platforms that are characterized by a very large continuity (i.e. several hundreds of kilometres across) of relatively shallow-carbonate facies. The development of facies models for this type of platforms is complex and the knowledge of the factors that controlled its sedimentary evolution is limited. This reflects the absence of recent analogues and the differences between the climatic conditions and oceanic circulation of the Holocene and the greenhouse intervals (e.g. Burchette & Wright 1992; Aurell et al. 1995; Wright & Burchette 1996). A significant difference between the recent rimmed flat-topped platforms and ancient epeiric carbonate platforms is the lateral development and extent of carbonate shoals. Recent rimmed flattopped platforms are limited in their progradation

by the surrounding deep ocean, whereas in epeiric platforms, the potential of progradation of the shoals over depositional slopes of few degrees results in a much higher potential of preservation of grain-supported facies (e.g. Droste 2006). The knowledge of the geometry and lateral continuity of these grain-supported facies and its relationship with the mud-supported, relatively deep facies is a key aspect for reservoir exploration and field development strategies. Successive epeiric carbonate platforms developed during the Late Jurassic in the eastern part of the Iberian Plate, in the so-called Iberian Basin. These platforms had low-angle depositional slopes separating the shallow and deep-water areas and have been interpreted as epeiric carbonate ramps (e.g. Aurell et al. 2003; Ba´denas & Aurell 2008). The high quality of exposure across the study area clearly shows the facies distribution across these carbonate ramps and thus enables the reconstruction

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 89– 111. DOI: 10.1144/SP329.5 0305-8719/10/$15.00 # The Geological Society of London 2010.

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of a transect showing the transition of facies from shallow to relatively deep sedimentation sites. This work summarizes the facies present, their distribution and the long-term sequence evolution of the Upper Jurassic epeiric carbonate platform in the northern Iberian Basin from the major transgressive event in the Middle Oxfordian (i.e. transversarium zone; Aurell et al. 2003) to the widespread mid-Berriasian regression. It provides a case study of all of the facies present and considers the sedimentary evolution of comparable carbonate ramp systems.

Methods and material The reconstruction of the sedimentary evolution of the Upper Jurassic epeiric carbonate ramps summarized here are based on an outcrop analysis across the exposures of the Iberian Chain of the NE Spain (Fig. 1). The data supporting the presented interpretations are derived from the PhD theses of the authors (Aurell 1990; Ba´denas 1999; Ramajo 2006). Other publications by the authors showing key aspects of

the sedimentological and sequence stratigraphy of the Upper Jurassic carbonate platforms of the Iberian Basin are listed in the reference section. A bed-by-bed analysis of 180 outcrop sections has been carried out over the last 20 years. A selection of key sections (1– 8 in Fig. 2) has been used to illustrate these interpretations. These sections are continuous, have undergone relatively little deformation and are between 150–250 m in thickness. Field observations include determination of lithology and facies, analysis of the discontinuity surfaces and identification of sedimentary structures. This information was completed by microfacies analysis of polished slabs and thin sections (around 40–60 slabs and 30 –40 thin sections per log). These data were used to determine the depositional environment and define the depositional sequences (from long-term to high-frequency cycles). The large lateral extent of the Upper Jurassic Iberian ramps (several hundreds of kilometres across) renders the correlation between sections of widely spaced localities difficult. Hence, the importance of the precise ammonite biozonation

Fig. 1. Palaeogeography of Western Europe (early Kimmeridgian), indicating the location of the Iberian Basin and surrounding uplifted areas. Modified from Dercourt et al. (1993).

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Fig. 2. Distribution of the Jurassic outcrops (areas in black) in NE Spain, showing the location of reference sections 1– 8.

(Oxfordian to lower Tithonian) established across the northern Iberian Chain that supports the age assignment and correlation between the logs located in the outer-ramp areas (e.g. Mele´ndez & Fontana 1993; Mele´ndez et al. 1997; Ramajo 2006). The correlation between shallow and deep ramp localities is based on other biostratigraphic data (such as lituolids and calcareous algae) and sequence stratigraphic principles. The now wellestablished method of using sequence stratigraphy in the study of ancient carbonate platforms (e.g. Loucks & Sarg 1993; Strasser et al. 1999; Pomar 2001) was used in a number of publications to correlate facies and sequences at different scales, from high-frequency cycles to long-term (i.e. second- and third-order) depositional sequences (e.g. Aurell & Mele´ndez 1993; Aurell et al. 2003; Aurell & Ba´denas 2004; Ramajo & Aurell 2008).

Geological setting and previous work During the Late Jurassic, shallow epeiric seas covered wide areas of western Europe. Terrigenous

and siliciclastic basins developed to the north, whereas carbonate sedimentation was dominant in the basins located around the deep oceanic areas of Tethys sea (e.g. Dercourt et al. 1993; Fig. 1). A large part of the central and western Iberian Plate was an uplifted high, the so-called Iberian Massif. Subsiding areas located eastwards in the Iberian Plate correspond to the Iberian Basin. This was an intra-cratonic basin developed during the Mesozoic extensional phase affecting the NE Iberian Plate (Salas & Casas 1993). Most of the Upper Jurassic sediments deposited in the northern Iberian Basin record a period of relative tectonic quiescence and broad regional subsidence (i.e. the Jurassic postrift stage). A period of more differentiated subsidence restricted to specific zones controlled by normal fault activity started at the latest Jurassic (i.e. the Early Cretaceous rifting episode; Salas & Casas 1993). During the Late Jurassic, the Iberian Basin was occupied by large east-facing carbonate ramps, with ammonite faunas showing a markedly Tethyan affinity. These ramps had low-angle depositional

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Fig. 3.

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slopes separating the shallow and deep-water areas. The Upper Jurassic platforms faced hurricanes and winter winds (Marsaglia & Klein 1983; Price et al. 1995) and deposition was dominated by storms (Ba´denas & Aurell 2001a). In contrast, the effect of tides was restricted. The fair weather wave base is thought to have been quite shallow, between 5–10 m, and was similar to the depth considered for ancient epeiric platforms and to that observed in comparable modern settings (e.g. Tucker & Wright 1990). The storm wave base is assumed to have been around 30– 60 m (Aurell et al. 1995, 1998). This figure is in the range of the effective storm wave base in the Persian Gulf (Immenhauser 2005). The lithostratigraphy of the Upper Jurassic of the Iberian Chain is summarized in Figure 3. Over the years, the interpretation of the distribution of the main lithostratigraphic units has evolved from a nearly ‘layer-cake’ stratigraphy (Giner 1980; Aurell 1990) to a more realistic model, showing the interfingering between the shallow platform facies and the marly and mud supported facies found in the outer-ramp localities (Fig. 3). The progress in the accuracy of the stratigraphic and sedimentological interpretation was favoured by a number of factors, including: (1) the increasing availability of biostratigraphic data, not only of the pelagic and hemipelagic fauna found in the open platform facies, but also of some benthic groups, such as lituolids and calcareous algae, which developed in shallow-water platform environment (Ba´denas 1999; Ramajo 2006); (2) the analysis of more closely spaced outcrops, filling the gaps between widely spaced sections, particularly in the transition from shallow to relatively deep ramp environments (e.g. Ba´denas & Aurell 2001b; Aurell & Ba´denas 2004; Ba´denas et al. 2008).

the sequence stratigraphy (e.g. Ba´denas & Aurell 2001a; Aurell et al. 2003; Ipas et al. 2005; Ba´denas et al. 2008; Ramajo & Aurell 2008) have integrated previous interpretations of the long-term sedimentary evolution of the platforms developed in the northern Iberian Basin in different hierarchical orders of sequences (Fig. 4): (1)

(2)

(3)

Sequence stratigraphy and sedimentary evolution The first sequence stratigraphic interpretation of the Upper Jurassic rocks of the Iberian Chain (Giner 1980; see Fig. 3) considered three long-term sequences, roughly corresponding to the Oxfordian, Kimmeridigan and Tithonian –Berriasian, respectively. Facies distribution and sedimentary evolution of these three sequences was further documented in Aurell & Mele´ndez (1993). More recent works on

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(4)

A transgressive –regressive super sequence, with a duration of c.15 Ma, that extends from the Oxfordian to the early Berriasian, with major deepening around the early –late Kimmeridgian boundary (Aurell et al. 2003). Five long-term, second- to third-order sequences, with a duration varying from 1.5 to 6.5 Ma, are distinguished, which have well-defined retrogradational-progradational facies evolution patterns (i.e. the Ox, Kim1, Kim2, Ti1 and Ti2 sequences in Fig. 4; Aurell et al. 2003). Sequence boundaries were placed at transgressive surfaces following stages of accommodation loss at the end of the sequences, with local emersion in shallow platform areas. The deposits formed coeval to relative sea-level fall (i.e. the forced regressive systems tract at the end of the sequence, sensu Hunt & Tucker 1993) have locally been identified as low-relief wedges developed in the shallow–deep ramp transition area of the Kimmeridgian sequences (Ba´denas & Aurell 2001a). Transgressive peaks or maximum flooding zones of these sequences are indicated by the maximum lateral extension of the open platform facies (Fig. 4). A variable number of depositional sequences with an average duration ranging between 0.5–1.5 Ma (see the Ti1-1 to Ti2-5 sequences in the Tithonian –lower Berriasian interval, Fig. 4) have been described for the Oxfordian (Ramajo & Aurell 2008), Kimmeridgian– lower Tithonian (Aurell & Ba´denas 2004; Ipas et al. 2005; Ba´denas et al. 2008) and middle Tithonian–lower Berriasian (Ba´denas et al. 2004). These classify as third-order sequences. High-frequency sequences equivalent to the medium and small-scale sequences of Strasser et al. (1999), have been related to orbital cycles influencing the sedimentation during the Oxfordian (Strasser et al. 2005),

Fig. 3. (Continued) Interpretation and distribution of lithostratigraphic units and long-term sequences of the Kimmeridgian– Tithonian of NE Iberia. (1) Giner (1980): an isochronous reefal-oolitic-oncolitic unit (the Higueruelas Fm, H) covering the outer ramp marls and micrites of the Sot de Chera Fm (SCh) and Loriguilla Fm (L); (2) Aurell & Mele´ndez (1993): the oolitic Pozuel Fm (P) partly equivalent to the outer ramp micrites (L); (3) Ba´denas (1999): an upper Kimmeridgian reefal unit (i.e. the Torrecilla Fm, T) grading laterally to the open marine Loriguilla Fm (L); (4) Ipas et al. (2005): Outer ramp micrites age-equivalent to the shallow oncolitic facies of the Higueruelas Fm (H).

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Fig. 4. Synthesis of the stratigraphy of the Upper Jurassic transgressive– regressive (T– R) cycle across the northern Iberian Ranges (see reference sections 1– 8 in Fig. 2) with age and main facies distribution of the five long-term depositional sequences defined in the Iberian Basin (modified from Aurell et al. 2003).

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Kimmeridgian (Ba´denas et al. 2003, 2005) and Tithonian –Berriasian (Ba´denas et al. 2004). The five long-term sequences indicated in (3) have particular facies patterns and are used below to describe the successive stages of sedimentary evolution of the Upper Jurassic Iberian platform. Field images of the distribution of these sequences are shown in Figures 5 and 6. The descriptions of the sequences will begin with an indication of the age and the main features associated to the key surfaces (boundaries and transgressive peaks), followed by the description of the facies distribution, from proximal to distal ramp environments. The presented long-term sequence stratigraphic models are also illustrated with representative examples showing the small-scale facies successions (high-frequency cycles) that form the building blocks. The long-term sedimentary evolution of the successive platforms is also summarized.

The Oxfordian sequence Boundaries and transgressive peak. The Oxfordian (Ox) sequence is located above the unconformity that developed around the Callovian –Oxfordian transition. This unconformity is the expression of an important regressive event at the end of the Callovian, which has been associated with widespread shallowing and local emersion of the open platform areas, causing a stratigraphic gap in the Late Callovian –early Middle Oxfordian (e.g. Ramajo & Aurell 2008). Marine sedimentation was fully recovered after the widespread flooding of the platform at the onset of the Middle Oxfordian transversarium zone (Fig. 4). This flooding continued through the bifurcatus zone, involving widespread development of spongiolithic facies across the northern Iberian Basin. The transgressive peak of the Oxfordian cycle can be interpreted to be located in different stratigraphic intervals depending on the location of the platform: either on top of the transversarium zone (Aurell & Mele´ndez 1993), at the top of the hypselum subzone (Ferna´ndez-Lo´pez & Mele´ndez 2004) or on top of the bifurcatus zone (Ramajo 2006). This last option, adopted in the interpretation of Ramajo & Aurell (2008), is supported by the overall facies distribution observed across the northern Iberian Basin (Fig. 4). After the maximum flooding stage, there was a progressive progradation of siliciclastic facies, with shallowing and a shallowing-upwards trend into the latest Oxfordian (Fig. 4). The upper boundary of the sequence developed in the mid-planula zone, involving a local stratigraphic gap, which may partly affect the hauffianum zone and the lower part of planula zone (Aurell et al. 2003).

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Facies distribution and high-frequency cycles. In the Oxfordian carbonate platform, there is a sharp transition between the shallow and the relatively deep facies (Fig. 7). The shallow platform areas were located to the NE, in the so-called Soria Seaway, and have been studied in detail in the Aldealpozo section by Alonso & Mas (1990), Pittet & Strasser (1998) and Strasser et al. (2005). Siliciclastic facies (i.e. sandstones with carbonate matrix) form the lower and upper part of the succession, whereas the middle part is dominated by a wide variety of wackestones to grainstones facies. Depositional environments in the shallow platform included lagoons with carbonate mud, peloids and oncoids, subtidal dunes with carbonate and/or siliciclastic grains, tidal-flats with algal-microbial mats and bird’s-eyes, and soils. A thick, marl-dominated wedge is locally found in the transition area between the shallow and relatively deep platform setting (Fig. 7). These marls contain sandy limestones (quartz, mica) with peloids and echinoderms. Basinwards, the open platform areas are usually represented by marl-limestone alternations that contain abundant sponges, tuberoids, and microbial crusts, as well as ammonites, belemnites, planktonic foraminifera and brachiopods (e.g. Ricla section in Fig. 8). Sponges (mainly Hexactinosan) are either preserved in the original upright position in wackestones, packstones, or boundstones (e.g. sponge bioherms in Ricla section 3, see Fig. 8), or they are toppled and broken in marly floatstones (Strasser et al. 2005; Ramajo 2006). Local development of swells in open platform areas is indicated by peloidal packstones with abundant glauconite as well as echinoderms, belemnites and ammonites (around section 5; see Figs 4, 7 & 8). The distribution of ammonite subzones shows highly variable sedimentation rates (compare de Ricla-Arin˜o logs in Fig. 8); sedimentary sequences are commonly condensed (Mele´ndez & Fontana 1993; Mele´ndez et al. 1997; Ferna´ndezLo´pez & Mele´ndez 2004; Ramajo 2006). High-frequency cycles have been identified across the different platform environments (Fig. 8). In shallow platform areas they consist of stacked metre-scale sequences that are commonly capped by tidal-flat facies or by palaeosols. In most cases, the facies evolution within these sequences displays a deepening, then shallowing trend. In the open platform environment, the equivalent sequences are much thinner (generally less than one metre) and are formed by bundles of beds bounded by ironencrusted surfaces. These sequences are interpreted as reflecting orbital eccentricity cycles (Fig. 8, Strasser et al. 2005). Sedimentary evolution. Two stages of evolution of the platform are distinguished in the facies

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Fig. 5. (a, b) Distribution of transgressive and regressive deposits of Kim1 sequence in section 3: transgressive deposits are characterized by silty lime mudstones and marls; highstand deposits (located above the dashed line in (b) include the progradation of oolitic, sandy dunes (with sets up to 4 m thick in (a) over marls with sandy tempestites. (c, d) Distribution of Ox, Kim1, Kim2, Ti1 and Ti2 sequences in section 4. Outer ramp lime mudstones and marls characterize the Kim1 and Kim2 sequences. Marly-dominated interval, found at the uppermost part of Kim1 sequence, is the coeval counterpart of the oolitic dunes of shallower sectors. Sharp facies change from outer ramp lime mudstones to middle-ramp oncolitic shoals represent the boundary between Kim2 and Ti1 sequences (see c).

distribution transect presented in Figure 7. In the lower stage 1 (from transversarium to midbimammatum zones) there is a sharp lateral change from mixed carbonate-siliciclastic shallow platform facies to condensed sponge and ammonite

limestones found in the open and relatively deep platform areas (Fig. 8). The observed aggradational geometry reflects lower rates of shallow-water carbonate production and limited offshore transport of the newly deposited sediments during a stage of

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Fig. 6. (a) Distribution of Ox and Kim1 sequences in section 5. Limestone with abundant sponges and tuberoids and peloidal-glauconitic facies characterize the Oxfordian sequence. (b) Outer ramp lime mudstones and marls of the Kim1 and Kim2 sequences in section 6: as in section 3, the boundary between Kim1 and Kim2 sequences is located above a marly interval. (c) Distribution of lime mudstones and marls of Kim1 and Kim2 sequences and the peloidal–bioclastic limestones of Ti1 sequence in section 5. Higher-order Ti1-1 and Ti1-2 sequences are indicated. (d) Distribution of the five higher-order sequences developed in Ti2 sequence in section 8.

98 M. AURELL ET AL. Fig. 7. (a) Isopach map and (b) palaeogeographic reconstruction of the Oxfordian platform. (c) Main facies distribution along two stages of sedimentary evolution. The location of reference sections 1– 6 is indicated.

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Fig. 8. Correlation between Middle Oxfordian high-frequency cycles defined in the interior and the outer platform areas (compiled from Strasser et al. 2005).

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rapid relative sea-level rise at the lower transgressive hemicycle and at the early highstand. During the upper stage, the latest Oxfordian, siliciclastic influence was much greater. In the proximal areas close to the Ebro Massif (section 2, see the palaeogeographic inset in Fig. 7b), a thick unit (up to 50– 70 m) of cross-bedded sandstones and conglomerates developed in the hauffianum-planula zone (Ramajo 2006). Quartz and mica also became abundant in the equivalent thick marly wedge developed in the transition area between the shallowand deep-water platform domains (see section 3 in Fig. 5). The unusual thickness of this marlydominated wedge reflects the existence of differential subsidence induced by the local tectonic development at the latest Oxfordian. Lowsedimentation rates and condensation are also evident in the open platform areas during this second stage of evolution. Sedimentation in the open platform is locally dominated by peloidalglauconitic facies (Figs 4 & 7). Glauconite suggests generally low sedimentation rates (Odin & Matter 1981), and iron crusts on some bed surfaces indicate intervals of prolonged absence of sedimentation.

The Kimmeridgian 1 and 2 sequences Boundaries and transgressive peaks. The Kimmeridgian strata consist of two depositional sequences that have a long-term transgressive-regressive facies evolution (Fig. 4). The Kim1 sequence embraces the latest Oxfordian to the middle part of the late Kimmeridgian (acanthicum zone). The transgressive interval of the Kim1 sequence developed during the early Kimmeridgian with the transgressive peak located in the upper divisum zone (Ba´denas & Aurell 2001a). The Kim2 sequence has the transgressive peak in the middle eudoxus zone and developed up to the end of the Kimmeridgian (Fig. 4). These two transgressive peaks are likely to reflect episodes of regional sea-level rise controlled by the regional tectonic development that affected several western European basins (Ba´denas & Aurell 2001a; Aurell et al. 2003). Facies distribution and high-frequency cycles. The overall facies and thickness distribution of the two Kimmeridgian platforms of the northern Iberian Chain is illustrated in the cross-section from shallow-water to relatively deep-water ramp realms (Fig. 9). Inner to middle-ramp oolitic and reefal facies dominated in the Basque–Cantabrian Basin, in the central parts of the Soria Seaway, and in the marginal areas of the Iberian Basin (Ba´denas & Aurell 2001a; Benito & Mas 2006). Siliciclastic input was significant in some marginal areas. The shallow-water reefal, skeletal, ooidal and siliciclastic facies graded basinwards into a

middle-ramp facies belt characterized by the existence of abundant skeletal and siliciclastic tempestites. Further offshore, there is a progressive transition from the tempestite facies belt to the rhythmic alternations of lime mudstones and marls with scarce fossils (ammonites, bivalves, brachiopods, sponges, benthic foraminifera) that characterized most of the outer-ramp setting (Figs 5d & 6b). The proportion of ammonites and siliceous (Dictyida and Lithistida) sponges increased towards the more distal areas. The three types of facies belts found in the inner-, mid- and outer-ramp environments are distinguished in varying shades in Figure 9. Facies are organized in high-frequency cycles, which are recorded from shallow-water platform to relatively deep-water ramp areas (Fig. 10). These cycles have been related to sea-level fluctuations and orbital cycles. In shallow-water areas highfrequency cycles may include a lower transgressive interval with pinnacle reef development, followed by the progradation of inner-ramp shoal and lagoon environments. These sequences were related to high-frequency sea-level fluctuations that affected the carbonate productivity (Aurell & Ba´denas 2004). Cyclical variation in shallow-water carbonate production led to changes of the amount of carbonate exported to deeper areas, and explains the origin of the high-order sequences recorded in middle –outer ramp settings (Ba´denas et al. 2003, 2005). Sedimentary evolution. The Kim1 sequence shows a long-term retrogradational-progradational pattern (Figs 4 & 9). The overall retrogradation of the outerramp lime mudstones over the more proximal marls has been related to a widespread relative sea-level rise during the early Kimmeridgian (Figs 4 & 9). The progressive deepening involved the local development of condensed successions with ammonites in the more open platform areas, represented by log 6 in Figure 9. The regressive interval of the Kim1 sequence is marked by the overall progradation of inner–middle ramp, oolitic-sandy facies (Fig. 5a, b). These shallow oolitic facies recorded in the Ricla marginal area (i.e. section 3, Fig. 5a) correspond to shoals with dunes 1–5 m high that migrated basinwards (to the SSW) under the influence of storm-induced return currents (Ba´denas & Aurell 2001b). Offshore to these facies there is a progressive decrease of sandy and oolitic tempestites in the well bedded lime mudstones and marls that characterize the outer-ramp setting. A rapid rise in sea-level at the onset of Kim2 sequence is shown in the marginal areas of the basin by a sharp facies change (Figs 4 & 9). In the Ricla area (section 3), metre-size coralgal patch reefs and micritic limestones with storm deposits sharply overlie the inner-ramp oolitic dunes of the

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Fig. 9. (a) Isopach map and (b) palaeogeographic reconstruction of the Kimmeridgian platform. (c) Main facies distribution of the two long-term transgressive-regressive sequences Kim1 and Kim2. The location of reference sections 1– 6 is indicated.

underlying sequence. This transgressive evolution is characterized by the aggradation of these reefal facies, previous to the rapid progradation of innerramp (lagoon) skeletal and silty limestones that took place at the upper part of the sequence (Ba´denas & Aurell 2001b). The offshore counterparts of these shallow facies are lime mudstones and marls, which contain abundant tempestites in more proximal areas (Fig. 10).

The Tithonian 1 and 2 sequences Boundaries and transgressive peaks. Differential tectonics involving the uplift of the marginal areas of the Iberian Basin, combined with a long-term regional sea-level fall, resulted in the progressive coastal offlap and progradation of the Tithonian carbonate platforms (Salas et al. 2001; Martı´nChivelet et al. 2002; Aurell et al. 2003). As a consequence, the western extension of the lower Tithonian carbonate platform is more reduced, and

they were not developed west to the area located around section 4 (see the palaeogeographic inset in Fig. 9). The Tithonian strata consist of two long-term depositional sequences bounded by a major discontinuity surface. The Tithonian 1 sequence (Ti1) is early Tithonian in age, with the transgressive peak in the lower hybonotum zone (Ipas et al. 2004; Ba´denas et al. 2008). The age of the Tithonian 2 sequence (Ti2), from middle Tithonian to early Berriasian was determined by the widespread presence of Anchispirocyclina lusitanica (Fig. 4). The transgressive peak is interpreted to have occurred around the middle part of the Tithonian (Ba´denas et al. 2004; Ipas et al. 2006). The Ti1 and Ti2 sequences have important differences in geometry and facies architecture, and will be described separately in the next sections. The overall thickness of Ti1 sequence slightly increases basinwards (from 50 –80 m), reflecting similar subsidence rates across the basin (Fig. 11).

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Fig. 10. Correlation between high-frequency cycles defined in the mid and the outer ramp areas in the lower part of sequence Kim2 (compiled from Ba´denas et al. 2005).

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Major change related to tectonic reactivation occurred at the onset of Ti2 sequence: important differential subsidence resulted in the formation of thick carbonate successions at the eastern Iberian Basin, out of the siliciclastic influence. Facies, sedimentary and sequential evolution in the Ti1 sequence. The facies associations found in the lower Tithonian carbonate ramp have been grouped in four types in Figure 11. A wide highenergy shoal characterized by oncolitic rudstones with scarce bioclasts is interpreted as the midramp environment, below fair weather wave base. The more inner parts of the platform are characterized by peloidal and bioclastic wackestones– packstones with calcareous algae, gastropods, bivalves, echinoderms, and by locally developed coral-microbial patch reefs that include also stromatoporoids and chaetetid sponges. The facies mosaic in these interior areas of the platform is complex, with the local development of high-energy oolitic shoals favouring the presence of restricted lagoons with sedimentation of mud-dominated facies (bivalve and gastropod wackestones) (Ipas et al. 2004). The outer-ramp lime mudstones consist of wellbedded micrites with scarce ammonites and benthic fossils (brachiopods, bivalves, sponges) similar to those found in most of the Kimmeridgian deposits. Basinwards, organic-rich laminated marls and lime mudstones were deposited during early Tithonian times (Rossi et al. 2001). Inshore, the outer-ramp lime mudstones grade into a facies zone characterized by the progressive gain of grain-supported facies, indicating its location around the storm wave base level. In this zone, there is a transition between the lime mudstones and the peloidal facies with scarce oncoids and bioclasts (bivalves and sponge spicules). The lower Tithonian Ti1 sequence is characterized by overall facies progradation. In particular, the evolution of the facies belt dominated by the oncolitic facies allows the identification of three higher-order sequences. These are defined by the existence of rapid progradational pulses (Ti1-1, Ti-2 and Ti1-3 sequences in Figs 4 & 11, see also Figs 5d & 6d). The basic building block of these sequences consist of metre-scale higher-order cycles (Fig. 12), which have been tentatively related to sea-level changes in tune with the short eccentricity cycle (Ipas et al. 2004; Ba´denas et al. 2008). Facies, sedimentary and sequential evolution in the he Ti2 sequence. During the development of the Ti2 sequence, the marginal zones of the basin were characterized by the large siliciclastic influence (see palaeogeographic inset in Fig. 11). Around

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the more western section 4, the siliciclastic facies are organized in thickening and coarsening-upward sequences reflecting the progradation of deltaic systems (Ipas et al. 2006). The offshore equivalent of these facies (i.e. sections 5, 6) include marlydominated intervals with some charophytes and ostracods, reflecting the episodic setting of freshwater environments. Away from this continental influence, tidal processes dominated in the interior areas of the carbonate platform, as indicated by the widespread presence of algal-laminated facies. These facies form the upper intervals of muddominated, metre-scale, shallowing-upward sequences. The high-energy facies developed in the mid-ramp environment are well represented further east (section 8). The facies are dominated by peloidal, oolitic and bioclastic packstones and grainstones. The Ti2 sequence has an unusual duration and thickness with respect to the previously defined sequences. It consists of five higher-order sequences (from Ti2-1 to Ti2-5 in Fig. 11; see also Fig. 6d), each one showing a well-defined transgressiveregressive evolution. All these five sequences are included in a single long-term depositional sequence because there are no major facies changes associated with the discontinuities bounding these higher-order sequences. These sequences have a range and sedimentary evolution comparable to the five third-order sequences 1–5 defined in Ba´denas et al. (2004) further east, in the nearby Montanejos and Salzedella sections. These two sections correspond to the more eastern and open domain of the Iberian Basin, showing the local development of calpionellid-rich mudstones around the maximum flooding of sequence 2 (see Fig. 11b for location). This flooding event is likely to be coeval with the transgressive maxima found in the middle part of sequence Ti2-2, which involved the occasional presence of marine facies over the marginal areas of the basin (i.e. section 4 in Figs 4 & 9). Particular types of shallowing-upward to symmetric (i.e. deepening-shallowing upward) highfrequency cycles were defined for each of the third-order sequences 1–5 in the MontanejosSalzedella area. These metre to decametre-scale sequences were related to the short-eccentricity orbital cycle (Ba´denas et al. 2004).

Discussion Comparison of the Upper Jurassic Iberian epeiric carbonate ramps with other age equivalent systems located in nearby basins has been used to consider the factors controlling the lithofacies character and distribution of the study area. Of particular relevance is the review of the more significant events

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Fig. 11. (a) Isopach map and (b) palaeogeographic reconstruction of the Tithonian platform. (c) Main facies distribution of the two long-term transgressive-regressive sequences Ti1 and Ti2, and respective higher-order sequences. The location of reference sections 4 –8 is indicated.

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Fig. 12. Correlation between high-frequency cycles defined in the lower part of the Ti1 sequence in some shallow platform localities (compiled from Ipas et al. 2004, 2005).

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affecting the western Tethyan basins during the Late Jurassic.

The Oxfordian carbonate platform During late Callovian– early middle Oxfordian, a global cooling event affected the Tethyan and peri-Tethyan basins and the deposits of epeiric seas became almost carbonate free around the Callovian–Oxfordian transition (e.g. Dromart et al. 2003; Ramajo & Aurell 2008). A widespread flooding event at the onset of the middle Oxfordian transversarium zone, involved the return of the open marine and carbonate sedimentation. This flooding event has been correlated with a ‘thermal maximum’ and warm conditions persisting throughout the late Oxfordian (e.g. Cecca et al. 2005). Sponge-dominated facies developed widely in the northern margin of the Tethys during the middle–late Oxfordian (see Olo´riz et al. 2003 and references therein). The flooding and recovery of marine carbonate sedimentation also involved the widespread development of spongiolithic facies in the open platform areas of the Iberian Basin. The dominance of sponges differs from those of other Tethyan and peri-Tethyan areas (Cecca et al. 2005); the middle Oxfordian of some western European basins was characterized by the widespread deposition of carbonates including coral reefs, whereas at the southern margin of the Tethys a maximum development of radiolarites occurred. There was general recovery of carbonate factories along the southern margin of Tethys from the end of the Oxfordian, with radiolarites diluted by periplatform oozes. Cecca et al. (2005) relate this palaeobiogeographic distribution to a latitudinal shift of the climatic (arid- humid) belts. The dominance of sponges and siliciclastics in the Iberian platform can be explained mostly by the sedimentation rates which were controlled by the long-term evolution of accommodation and tectonics. The sponge fauna of the Iberian Basin is dominated by low-diversity Hexactinosan sponges, that reflect very low sedimentation rates and nutrient levels in the outer platform areas (Krautter 1998). This is coherent with the long-term sea-level rise around the middle–late Oxfordian transition, during which the siliciclastic and carbonate sedimentation rate decreased (see stage 1 in Fig. 8). However, in stage 2, decreasing accommodation rates led to progradation of siliciclastics from the hinterland and increasing sediment input to the outer platform, which impeded the colonization of the seabed by sponges. A similar interpretation of accommodation-controlled distribution of spongiolithic facies is given in Olo´riz et al. (2003) in the Oxfordian facies of southern Spain (see also Leinfelder et al. 1994).

The absence of coral reefs in the shallow areas of the Oxfordian platform (e.g. Aldealpozo section; Fig. 8) can be attributed to eutrophication due to high siliciclastic input or, more probably, to low temperatures due to cool waters entering throughout the Soria Seaway from the North Atlantic Ocean (Pittet 1996). As indicated by Cecca et al. (2005), a detailed palaeoecological study of the Oxfordian reefs to confirm the controlling effect of sea-water temperatures and/or latitudinal shift of climatic belts on the distribution of sponges and coral-reefs in the Oxfordian platforms, is still lacking.

The Kimmeridgian carbonate platform The Oxfordian–Kimmeridgian transition represents a sharp change in facies and sedimentary conditions over the Iberian carbonate platforms. Shallow areas of the Kimmeridgian carbonate ramp located away from the siliciclastic influence were characterized by oolitic sand dunes and the growth of coralmicrobial reefs. Sponges became restricted to certain open platform localities, and these areas were mostly dominated by rapid deposition of carbonate mud and marls. The facies change from condensed, spongedominated facies (Oxfordian) to thick accumulations of carbonate mud (Kimmeridgian), indicates a general trend to more optimum conditions for coral reef growth and carbonate deposition that could be related to the combined effect of (1) increasing temperatures, coherent with a general warming through the Oxfordian–Kimmeridgian (Gro¨cke et al. 2003), and (2) decreasing siliciclastic input, due to progressive tectonic quiescence throughout the Kimmeridgian, and also probably, to less humid climatic conditions in the hinterland. In the Jura Platform, a decrease in siliciclastic influx during the Kimmeridgian –Tithonian, compared to the Oxfordian, has also been inferred as an evolution to a less humid climate (Rameil 2005). Extensive growth of carbonate-producing organisms occurred along the northern Tethys shelf during the Kimmeridgian. Shallow epicontinental seas covered wide areas of the western Tethys (see Fig. 2), with the development of carbonate platforms such as the Jura Platform (Colombie´ & Strasser 2005; Jank et al. 2006; Strasser 2007; Colombie´ & Rameil 2007), the southern and western edges of the SE Basin in France (Seguret et al. 2001) and the platform developed south of the Armorican Massif (Olivier et al. 2008). These were slightly sloping platforms bounded by more or less continuous bioclastic and oolitic sand barriers and corals reefs. The Iberian carbonate ramp had a particular style in the distribution of the shallow facies belts, characterized by a limited extent of lagoons and

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the absence of peritidal facies. This distribution is in contrast to the Jura Platform, for example, where the interior lagoon and peritidal facies have a widespread development. The observed differences could be related to their particular palaeogeographical location controlling the overall platform morphology. The Jura Platform developed around the emerged areas of an island archipelago and had a flat morphology, which favoured the presence of wide lagoons and tidal-flats in the interior platform areas. In contrast, the Iberian ramp developed along the margin of the Iberian Massif, and faced the Tethys Ocean, thus probably receiving a much higher frequency of storms (Ba´denas & Aurell 2001a). As a consequence, protected shallow areas (lagoons and tidal-flats) consist of discontinuous facies belts with a more reduced lateral extension. Despite the differences in type and distribution of shallower facies, the outer-ramp areas of the Jura and the Iberian platforms were similar and were characterized by deposition of lime mudstones and marls partly originated by resedimentation of shallowwater carbonates (Bartolini et al. 2003; Ba´denas et al. 2003, 2005; Colombie´ & Strasser 2003).

The Tithonian carbonate platform Compared to the Kimmeridgian, temperatures were cooler during the early Tithonian. This was followed

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by a gradual increase in temperature across the Jurassic –Cretaceous boundary (Gro¨cke et al. 2003). The shallow-water areas of the Iberian carbonate ramps saw a change to a new facies-style at the Kimmeridgian –Tithonian transition, from oolitic shoal and coral reef development to the widespread presence of oncolitic rudstones at the early Tithonian (Ti1 sequence). In the northern part of the Iberian Basin, oncolitic facies form most of the high-energy, mid-ramp facies belt (Fig. 11). To the south, however, the oncolitic facies distribution is more patchy, with frequent lateral changes into oolitic and reefal facies (Ba´denas et al. 2008). The development of wide, but at the basin-scale laterally-discontinuous, oncolitic belts at the initial stages of the Tithonian may be related to local factors rather than global ones. The combination of tectonic uplift and sea-level fall at the Kimmeridgian –Tithonian transition may have caused a change in the platform morphology, favouring oncoid growth in areas with relatively low-energy conditions and lower background sedimentation rates. The abundance of oncoids reflects the water transparency and trophic levels, which are driven by the variable influx of terrigenous sediment onto the shallow carbonate platform (Ve´drine et al. 2007). Significant siliciclastic influence occurred in certain areas of Kimmeridgian ramp. In contrast, the shallow areas of the Tithonian ramp were almost

Fig. 13. Overall distribution of the grain-supported carbonate facies (dark grey), coarse, siliciclastic-dominated facies (light grey) and mud-supported facies (white); compiled from Figures 7, 9 and 11.

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free of siliciclastic sediments, and pure carbonate sediments were deposited in the lagoonal areas. Terrigenous influx was low, mainly due to the presence in the inland of tectonically-controlled depressions that acted as traps for siliciclastic sediments (i.e. the newly formed Cameros Basin; Martı´n-Chivelet et al. 2003). In consequence, the clearer water and lower values of nutrients favoured the growth of oncoids during the early Tithonian. The carbonate platforms developed from midTithonian–early Berriasian were restricted to the eastern part of the Iberian Basin, and the outcrop control is more limited compared to the previous stages of ramp development. The observed vertical and lateral facies distribution indicates a major change of depositional style of the platform, with a more complex relationship between the different shallow facies. These platforms are characterized by a mosaic of tidal-flats, protected lagoons, and peloidal-bioclastic-oolitic shoals, instead of the well-defined facies belts of the Kimmeridgian ramps. Therefore, these platforms are more similar to the ‘Jura Platform style’.

Synthesis A key aspect of this case study is the identification of the lateral and vertical relationship between the shallow- and relatively deeper-water facies. The evolution of the sedimentary systems is summarized in Figure 13. The depositional system is the result of a combination of different local, regional and global factors, including: (1) tectonism, controlling the differential subsidence, the overall basin shape and the amount of the siliciclastic supply; (2) the balance between changes in accommodation and carbonate production (depending on the ecological potential of the successive Upper Jurassic platforms), which determines the amount of progradation of the carbonate grain-supported facies (see Fig. 13); and (3) the volume of sediment transported offshore by storm-induced flows, which tends to reduce the inclination angle of the depositional slopes favouring the potential of progradation of the carbonate shoals. This fact, combined with relatively low rates of creation of accommodation space, may explain the greater preservation of grainsupported facies in the upper regressive part of the Kimmeridgian and lower Tithonian sequences. Further comparative studies with other ageequivalent platforms, will help to improve the understanding of the interaction between the factors controlling the development of the two main types of Upper Jurassic carbonate platforms described in the Iberian Basin: either low-angle carbonate ramps with ‘laterally extensive mid-ramp facies belts’ (Kimmeridgian, lower Tithonian), or platforms characterized by a less predictable ‘facies mosaic’,

as observed in the shallow areas of the Oxfordian and middle –upper Tithonian platforms.

Conclusions The Upper Jurassic– lowermost Cretaceous succession recorded in the northern Iberian Basin consists of one super sequence, which has a long-term transgressive-regressive evolution: from the midOxfordian to mid-Kimmeridgian long-term transgression, to the progressive offlap and regression until the mid-Berriasian. The Upper Jurassic–lowermost Cretaceous super sequence consists of five long-term, secondto third-order sequences (ranging in duration from 1.5–6.5 Ma) that have well-defined retrogradational-progradational facies evolution patterns. These five depositional sequences differ in overall depositional geometry and ecology. The Oxfordian sequence shows a sharp transition, from shallow siliciclastic-carbonate (ooidal, skeletal, peloidal) facies to condensed (i.e. spongiolithic, peloidal, glauconitic) facies in the open platform domain. Coral-microbial reefs and oolitic-peloidal-skeletal shoals characterized the shallow areas of the two Kimmeridgian sequences. The early Tithonian sequence is characterized by shallow oncolitic-peloidal and skeletal facies. Thick lime mudstone successions (i.e. rhythmic marls-mudstone alternations) are found in the open platform areas of the Kimmeridgian–lower Tithonian carbonate ramps. The middle Tithonian to lower Berriasian platform is characterized by a thick succession with metrescale shallowing-upward sequences with local development of peritidal, algal-laminated caps. The facies distribution and long-term sequence evolution of the Upper Jurassic epeiric carbonate ramps of the northern Iberian Basin summarized in this work, may potentially be useful in the reconstruction of the overall facies distribution in less accessible or buried epeiric carbonate platforms. The differences observed in the successive Upper Jurassic platforms are the result of the interplay between tectonics (major episodes of differential subsidence occurred around the Oxfordian – Kimmeridgian transition and onwards from the midTithonian), regional changes in sea level (with major transgressive events at the middle Oxfordian, early Kimmeridgian, and mid –late Kimmeridgian), the ecological potential of the platform (controlling the amount of the carbonate produced in shallow areas), and the volume of the offshore transported sediment by storm-induced flows, which is determinant in the potential of progradation of the carbonate shoals. The authors are indebted to reviewers F. S. P. van Buchem and H. Hillga¨rtner, who helped to improve the original

JURASSIC CARBONATE RAMPS IN SPAIN versions of the manuscript. Financial support was provided by the Spanish and Arago´n Governments (Projects CGL2007-62469/BTE and CGL2008-1237/BTE and Grupo Consolidado ‘Reconstrucciones Paleoambientales).

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Sequence stratigraphy and carbon isotope stratigraphy of an Aptian mixed carbonate-siliciclastic platform to basin transition (Galve sub-basin, NE Spain) J.-C. EMBRY1,2,3*, E. VENNIN1,4, F. S. P. VAN BUCHEM2,5, R. SCHROEDER6, C. PIERRE7 & M. AURELL8 1

Muse´e National d’Histoire Naturelle, 43 rue Buffon, 75005, Paris, France

2

Institut Franc¸ais du Pe´trole, Geology-Geochemistry Division, 1-4 avenue de Bois Pre´au, 92852 Rueil-Malmaison, France 3

Present address: Statoil ASA, Research Centre Bergen, Geology, 90 Sandsliveien, Bergen 5020, Norway

4

Present address: UMR CNRS/uB 5561 Bioge´osciences, Universite´ de Bourgogne, 6 bd Gabriel, 21000 Dijon, France 5

Present address: Maersk Oil Qatar AS, P.O. Box 22050 Doha, Qatar

6

Senckenberg Museum, 25 Senckenberganlage, D-60325, Frankfurt/Main, Germany 7

Universite´ Pierre et Marie Curie, 4 place Jussieu, 75252 Paris, France 8

Universidad de Zaragoza, 50009 Zaragoza, Spain

*Corresponding author (e-mail: [email protected]) Abstract: A 380 m thick Aptian platform to basin transition has been studied along a 16 km long transect of excellent and continuous outcrops in NE Spain. The series has been dated using biostratigraphy (foraminifera and ammonites) and carbon-isotope stratigraphy, and has been subdivided at four scales of depositional sequences. The Aptian marine succession is subdivided into two-large scale sequences separated by a middle Aptian sub-aerial exposure surface. A characteristic trend of the floral-faunal fossil assemblages is present, which evolves from orbitolinid-ooid dominated ramps in Sequence I-1, to a coral-stromatoporoid-microbialite dominated platform in Sequence I-2, to a rudist-dominated platform top in Sequence II-1, and finally to a second episode of orbitolinid-ooid dominated ramp system in Sequence II-2. There was an influx of siliciclastic sediments at the base and at the top of this succession. The detailed carbon-isotope curve measured along the Miravete section and covering almost the complete Aptian succession, is compared with published Aptian curves recorded in both basinal and carbonate platform settings along the northern and southern NeoTethys margins. It shows that the Galve sub-basin curve represents all the major isotope excursions of the lower and upper Aptian, in a dominantly shallow-water succession.

Intensive research during recent decades has demonstrated that Barremian –Aptian shallow water and pelagic sedimentation in the NeoTethys realm was strongly influenced by climatic and environmental changes affecting ecosystems and the geochemistry of the oceans (e.g. Weissert et al. 1985; Weissert & Breheret 1991; Bralower et al. 1994, 1999; Fo¨llmi et al. 1994; Lini 1994; Jenkyns 1995; Erbacher et al. 1996; Gro¨tsch et al. 1998; Menegatti et al. 1998; Moullade et al. 1998; Erba et al. 1999; Masse et al. 1999). An added interest of the Barremian– Aptian interval is that it also represents a significant petroleum system, which is particularly prolific in the Middle East (e.g. Murris 1981; Vahrenkamp

1996; van Buchem et al. 2002; Yose et al. 2006; Droste 2009). Although the climate of the Cretaceous is considered to have been relatively equable, and the warmest of the Phanerozoic (Hallam 1984), there is now a persistent stream of publications suggesting the existence of cooler times (Kemper 1987; Walter 1991; Weissert & Lini 1991; Price 2000; Hochuli et al. 1999), including in the late Aptian (Puce´at et al. 2003; Steuber et al. 2005). The early Aptian, in contrast, is thought to have had a more humid climate, due to a higher volcanic and tectonic activity associated with high rates of production of oceanic crust and increasing volcanic activity

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Engineering Geology Special Publications, 329, 113–143. DOI: 10.1144/SP329.6 0305-8719/10/$15.00 # The Geological Society of London 2010.

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on the Ontong Java Plateau. It was also the time of the first widespread deposition of organic-rich black shales of the Oceanic Anoxic Event (OAE) 1a (Weissert et al. 1985; Ruffel & Batten 1990; Larson 1991; Weissert & Breheret 1991; Weissert & Lini 1991; Bralower et al. 1994; Fo¨llmi et al. 1994; Lini 1994; Jenkyns 1995; Erbacher et al. 1996; Menegatti et al. 1998; Mutterlose & Bo¨ckel 1998; Bralower et al. 1999; Ruffel & Worden 2000; Weissert & Erba 2004). A limitation of many of the cited studies is that they cover only part of the Aptian succession. The shallow-water carbonate systems along the northern NeoTethys margins, for instance, commonly terminate in the lower part of the lower Aptian, before the Deshayesites deshayesi ammonite zone (Funk & Briegel 1979; Clavel et al. 1986; Funk et al. 1993; Fo¨llmi et al. 1994; Arnaud et al. 1998; Wissler & Weissert 2002; Wissler et al. 2003; Embry 2005; Fo¨llmi et al. 2006). Along the southern NeoTethys margin, on the Arabian plate, these platforms continued to exist up to the lower part of the late Aptian (Hillgartner et al. 2003; Masse et al. 1999; Pittet et al. 2002; van Buchem et al. 2002, 2009; Yose et al. 2006). However, in a few locations in the Mediterranean area, upper Aptian platforms are reported, such as in Greece (Gavrovo platform in Grotsch et al. 1998), Croatia (Husinec & Jelaska 2006), and the Galve sub-basin in NE Spain (Vennin & Aurell 2001; Bover-Arnal et al. 2009). The focus of this paper is on the Galve sub-basin which has continuous exposure of the platform to basin transition (380 m thick, 16 km long), covering both the lower and upper Aptian, and is thus well suited to study the Aptian depositional sequences and facies evolution. This paper builds on previous work on these outcrops by Vennin & Aurell (2001), and integrates sedimentology, palaeontology, palaeoecology, sequence stratigraphy and carbon isotope stratigraphy. Compared to the previous work it presents new biostratigraphic information (orbitolinids), more detailed sedimentological data, and a complete carbon and oxygen stable isotope curve for the Aptian. The proposed sequence stratigraphic model and the carbon isotope curves are compared to other time-equivalent successions in the northern NeoTethys and the Middle East.

Methods and materials This study is based on five continuous and wellexposed outcrops (Aliaga, Peral, Miravete, Portoles and Villarroya de los Pinares) (Figs 1 & 2), located in the eastern part of the Galve sub-basin along the eastern side of the Miravete anticline. These sections were logged in detail and sampled. A total of 250 samples were collected and used for microfacies analysis and biostratigraphic determinations based mainly on their orbitolinid content

(R. Schroeder, Fig. 3). A semi-quantitative analysis of the orbitolinid shape, size and abundance, combining field observations and thin section analysis, was also performed in order to study their morphological response to the rapid variations of the palaeoenvironmental conditions. In addition, sampling with a tight grid at an average spacing of 1 m was carried out along the Miravete section to measure the d13C and d18O curves. Because it was not possible to separate macro- or microfossils from the carbonate matrix, the carbon and oxygen isotope analyses were performed on the bulk CaCO3. The 106 collected samples were powdered and analysed with X-rays in order to detect the possible presence of dolomite which may interfere with the original isotopic signal. The samples were preferentially drilled in the micritic matrix and fine grained lithologies. In addition, to avoid biological fractioning or ‘vital effect’ on d13C and d18O values in the microbialcoral boundstones, the samples were carefully drilled in micritic matrix excluding the microbial encrustings. The samples were reacted with 100% phosphoric acid at 25 8C, and the resulting CO2 gas was analysed using a Finnigan Delta E triple collector mass spectrometer from the Universite´ Pierre et Marie Curie, Paris. The isotopic compositions are expressed in the conventional notation relative to the Peedee belemnite (PDB) reference. The analytical precision is 0.01% for both d18O and d13C measurements. The sedimentological and palaeoenvironmental interpretations are based on standard microfacies and facies analysis. The trends of increase or decrease in the accommodation/sediment supply ratio are determined on the basis of the palaeobathymetric interpretations, the preservation of sedimentary structures, the stacking pattern, and significant stratigraphic surfaces. The final building of the 2D sequence stratigraphic correlation model is constrained by the hierarchy of the depositional sequences defined for each outcrop section and the time control provided by the integration of biostratigraphic and stable isotope chemostratigraphic data. The analysis of aerial photographs was also used to control the correlation and the large-scale depositional geometries of the carbonate geobodies. Similar approaches can be found in the literature (e.g. van Wagoner et al. 1988; Homewood et al. 1992; Read 1995; Homewood & Eberli 2000; Pittet et al. 2002; van Buchem et al. 2002).

Geological settings Structural settings The Mesozoic Iberian trough located in the northeastern part of the Iberian craton is an extensional basin resulting from the progressive opening of the

CRETACEOUS PLATFORM MARGIN IN SPAIN 115

Fig. 1. Location and palaeogeographic reconstruction of the Galve sub-basin (study area) and the six other sub-basins (Oliete, las Parras, Morella, El Perello´, Salzedella and Pen˜agolosa) in the eastern Iberian Maestrat basin during the Lower Cretaceous (modified from Salas & Guimera 1996). These seven well-differentiated subsiding sub-basins are formed during the Early Cretaceous rift-stage.

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Fig. 2. Geological settings of the Galve sub-basin: (a) simplified geological map and location of the five studied outcrop sections. (b) Main Lower Cretaceous lithostratigraphic units (see text for detail).

Central Atlantic Ocean and the western NeoTethys (Salas & Casas 1993; Salas et al. 1997). The development of this basin is marked by two major phases of rifting from the late Permian to Trias and from the late Oxfordian to middle Albian. The second phase of rifting results in the progressive resetting of the Jurassic carbonate platform system and the

formation of new NW–SE trending basins such as the Maestrat Basin (Salas & Guimera 1996). The displacement along the NNW –SSE and ENE – WSW syn-sedimentary faults during this extensional episode controlled the thickness of the Cretaceous deposits and led to the subdivision of this basin into seven sub-basins: Oliete, las

CRETACEOUS PLATFORM MARGIN IN SPAIN Fig. 3. Villarroya de Los Pinares, Miravete and Aliaga outcrop sections: lithostratigraphic units, stratigraphic position of the collected ammonites and orbitolinids determined by Schroeder for this study and calibration against the previous orbitolinid biozonation proposed by R. Schroeder (1964). The proposed ages are based on the new biostratigraphic analysis. They fit with the orbitolinid biozonation established by Schroeder and the datings previously proposed at a regional scale (Canerot 1974; Murat 1983; Martinez et al. 1994). 117

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Parras, Galve, Morella, El Perello`, Salzedella and Pen˜agolosa (Fig. 1) (Salas & Casas 1993; Salas & Guimera 1996; Soria et al. 1998; Salas et al. 2001; Liesa et al. 2006). These basins were opened towards the ESE along the northern Tethyan margin. At the present day, they show a reactivation of the early Cretaceous NNW–SSE and ENE –WSW extensional faults associated with Alpine compression: close to the studied outcrops in the Galve sub-basin, one of the best examples is the NNW– SSE trending Miravete fault acting as a normal fault during the Cretaceous rifting episode and later reactivated as a reverse fault during Alpine compression generating a thrusting through an anticlinal structure known as the ‘Aliaga-Miravete anticline’ (Fig. 2). The hinge of this faulted anticline disappears under the thrust between Aliaga and Miravete localities (Fig. 2). According to the previous facies and structural analyses of the Cretaceous deposits, the Galve sub-basin which is the focus of this present study, shows a proximal-distal polarity oriented from the NW to the SE and its Cretaceous palaeogeometry and depositional thicknesses were largely controlled by the ENE – WSW-trending listric normal faults and the steeper NNW –SSE normal faults producing small half-graben structures tilted towards the SE (Figs 1 & 2) (Vennin & Aurell 2001; Liesa et al. 2006).

Lithostratigraphic framework Previous studies of Lower Cretaceous deposits in the Galve sub-basin (Schroeder 1964; Gautier 1980; Canerot et al. 1982; Martin-Closas 1989; Salas et al. 1995; Soria 1996; Simon et al. 1999) proposed a subdivision of eight unconformity bounded lithostratigraphic units (Fig. 2). The first three units are characterized by significant variations of facies and thickness (up to 600 m). The Castellar Formation, of Hauterivian to early Barremian age, is made up of lutites which evolve to marls and lacustrine limestones rich in charophytes. The Camarillas Formation, of early Barremian age, is mainly composed of sandy and lutitic fluvial deposits. The Castellar and Camarillas formations correspond to the continental ‘Wealdian facies’. The Artoles Formation, of late Barremian –early Aptian age, consists of alternations of green marls and bioclastic limestones rich in benthic foraminifera (miliolids and orbitolinids) and oysters (Canerot 1974; Vennin et al. 1993; Vennin & Aurell 2001). The vertical facies evolution records a progressive transition from a continental to a marine environment. The Aptian deposits are divided into three formations (Fig. 2): the calcareous Chert Formation consisting of benthic foraminifers, ooids and green algal-rich limestones passing laterally and vertically into the Morella Formation (composed of marls and

bioclastic limestones rich in oysters and terrigenous siliciclastics; Canerot 1974; Vennin & Aurell 2001). This marly formation does not occur in the Galve sub-basin. The subsequent Forcall Formation, composed of massive marls and intercalations of limestones, is overlain by the calcareous Villarroya de los Pinares Formation. The marine units reach a thickness of up to 400 m in the sector of Villarroya de los Pinares. The first lower Aptian deposits of the Chert Formation are dated with benthic foraminifers such as Choffatella decipiens and Praeorbitolina sp. (Schroeder 1964; Canerot et al. 1982). However, a Barremian age for the base of this formation is not excluded (Canerot 1974). The combination of the ammonites (Dehayesites deshayesi and Dufrenoyia furcata) and the orbitolinids (abundant Palorbitolina lenticularis) discovered in the Forcall Formation clearly indicate an early Aptian age (Schroeder 1964; Canerot 1974; Murat 1983; Martinez et al. 1994). The presence of Mesorbitolina parva and Pseudochofatella cuvillieri in the Villarroya de los Pinares Formation confirms an early late Aptian age (Fig. 3; Schroeder 1964; Canerot 1974; Canerot et al. 1982; Vennin & Aurell 2001). These Urgonian carbonate deposits are overlain by three upper Aptian –Albian formations. The Benasal Formation, composed mainly of marly facies and bioclastic limestones with intercalated lignitic layers (Canerot et al. 1982), passes laterally northwards into the Escucha Formation (characterized by continental deposits). The overlying Utrillas Formation consists of continental deposits including sandstones and clays (Canerot 1974; Canerot et al. 1982). The absence of marine fossils, the presence of large silicified tree trunks, the abundance of kaolinite and iron attests of the continental origin of the deposits. In this study, a refinement of the time framework for the upper Barremian–Aptian deposits is proposed, based on new orbitolinid data, carbon isotope chemostratigraphy and sequence stratigraphic considerations.

Facies analysis and depositional environment Twelve facies types (26 sub-facies) have been distinguished, arranged into five depositional environments (Table 1). These are, from proximal to distal: (1) continental to coastal domain; (2) inner platform; (3) Oolitic/orbitolinid sand shoal; (4) Coral/stromatoporoid/microbial reef; and (5) outer platform (lower offshore). The important vertical and lateral facies changes are captured in three depositional models (Fig. 4) which are typical for specific phases of the evolution of this sedimentary system: (a) an oolitic-orbitolinid,

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Table 1. Facies classification and interpretation for the Aptian deposits of Galve sub-basin (legend in Fig. 4).

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Fig. 4. Depositional models of the Aptian formations of the Galve sub-basin. (a) Chert and Benasal Formation: oolitic-orbitolinid mixed carbonate-siliciclastic homoclinal ramp; (b) Forcall and lower part of Villarroya de los Pinares Formation: microbial-coral rimmed shelf; (c) Upper part of Villarroya de los Pinares Formation: rudist mud flat; (for legend see Fig. 4).

mixed carbonate-siliciclastic ramp; (b) a rimmed shelf dominated by corals and microbialites; and (c) a rudist mud flat. Continental to coastal plain environment. This environment groups facies F1 to F3 together (Table 1). Facies F2 consists of marly deposits, rich in plant remains and pedogenic horizons (F1) in which channelized sandy mica-bearing deposits are intercalated. Marls, rich in pedogenetized

horizon, and lignite layers are typical of low energy and form in the flood plain (F1). This facies is rich in well-preserved freshwater fauna and flora, including charophyte oogonia, bivalves, gastropods, and ostracods. The sandy deposits (F3a) show oblique crossbedding and parallel laminations, fining-upward sequences and erosive bases. Locally fine- to coarsegrained sandstones are channelized and form lateral accretion bars. These coarse to fine sands are locally

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mica or glauconite-rich. This siliciclastic facies is interpreted as a continental, high energy setting of the fluviatile environment. The local presence in the sandstones of gastropods, bivalves, green algae and foraminifers of marine origin (F3b) is interpreted as evidence for occasional marine influence in the coastal plain setting (Table 1; Fig. 5a). Inner platform. This environment is characterized by facies with a high carbonate mud content, and a variable siliciclastic and bioclastic content (F4 to F6, Table 1). The faunal associations are dominated by rudists, gastropods, miliolids and ostracods, and are typical for a restricted to semi-restricted environment. Several sub-environments may be distinguished, representing permanently subaqueous to periodically exposed areas: Inner platform bioherms (Fig. 4b): accumulations of well-preserved rudists, gastropods, corals and pellets with geopetal fabric (F5c and F6a, b, Table 1 and Fig. 5b, c) in a micrite-rich sediment commonly form in a subtidal lagoon protected by a reef belt (see below). Inner platform-mud flat (Fig. 4c): this environment is characterized by an alternation of rudistdominated packstones to floatstones (F6a, b; Fig. 6a, d) and wackestones –packstones rich in miliolids, oysters and gastropods (F4a, b to F5a; Fig. 5b). These rudist bivalves (Fig. 5c), mainly Toucasia sp., are characteristic of muddy, quiet water, inner platform environments and form tabular accumulations (Masse & Philip 1981). The homogeneity of the facies, the low lateral thickness variations of the rudist bio-accumulations and the high abundance of peloids, miliolids and gastropods indicate a homogeneous shallow subtidal flat depositional environment. The mud flat environment is periodically emergent in the intertidal to supratidal zone, resulting in polygonal cracks or bird’s-eyes filled by vadose silt and cements (Fig. 4a). Oolitic/orbitolinid sand shoal. This facies association consists of packstones, grainstones and floatstones organized in shoals, either dominated by ooids, fragmented bioclasts, orbitolinids or siliciclastic sands (F3b to F10a, Table 1). The crosslaminated foresets observed in the sandy deposits indicate a terrigenous supply from east to west parallel to the palaeocoast line (Fig. 5a). These sandstones comprise plane to oblique stratifications, erosive bases. They show strong lateral facies variations and are interpreted as a mixture of low and high energy facies. The oolitic shoal facies is rich in benthic foraminifers (F7a, b and c; Fig. 5d), and the floatstones and packstones are rich in gastropods, miliolids and

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orbitolinids (F4b, F5a and b and F10a and b). The gradual lateral evolution of facies in a proximal (ooid-rich; Fig. 5d, e) to distal (orbitolinid-rich; Fig. 5f, g) orientation and the common occurrence of storm-influenced deposits (F5b and c) indicate a good communication with an open marine environment compatible with a homoclinal ramp depositional setting (Ahr 1985; Burchette & Wright 1992). Grain composition, consisting of ooid type, bioclasts and aggregates, is related to type and intensity of sediment source, hydrodynamics and depositional topography of the shoal (Loreau 1982). Low-dipping shoals (108) or tabular beds with a large surf zone promote the formation of highly bioturbated facies rich in concentric ooids, aggregates and more rarely fibro-radiate ooids (F7b, c, and d). This environment is interpreted as a moderate to high energy setting with permanent water agitation (shoreface to foreshore). Coral/stromatoporoid/microbialite reefs. This environment is characterized by the presence of corals, stromatoporoids, rudists and microbialites (Thaumatoporella, Bacinella, Lithocodium and Koskinobullina), which may form buildups features of different dimensions (Fig. 6). It groups five subfacies (F8a, b and F9a, c) corresponding to bioclastic limestones with variable grain-size and build-up facies. The main components are corals (Fig. 5e), microbialites and stromatoporoids floating in micritic matrix. The dominant coral framework is associated with abundant red algae (encrusting), microbialites and is encrusted by serpulid-worm tubes. These facies present an aggrading and then prograding basinwards geometry. Within the coral-stromatoporoid facies, two reef-growth phases are observed (Fig. 6c): (1) initially, a phase of vertical accretion with plurimetric thick bioherms rich in branching and solitary corals and microbialites (Bacinella, Thaumatoporella, Lithocodium and Koskinobullina) (F9b; Fig. 6b, e and f ). The light-dependent microbial association within this coral-microbialite reef indicates clear and shallow waters with oligotrophic to mesotrophic conditions (Leinfelder et al. 1993; Dupraz & Strasser 1999, 2002). The base of the bioconstructions comprise packstones with orbitolinids (F10a) and provide a firm substrate for the establishment of the reef builders. (2) In the second stage, these bioconstructions spread out laterally to form biostromes pluri-kilometric in extension and only a few metres thick. These biostromes show a more significant amount of colonial corals and stromatoporoids (F9c; Fig. 6e, f and g) and a decreasing content in encrusting microbialites (Bacinella, Thaumatoporella and Lithocodium). Corals are commonly affected by Lithophaga and the reef facies display an intense bioerosion and

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evolve into wackestones rich in rudists and miliolids (F6a) interpreted as restricted lagoonal deposits. The bioerosion weakened the coral framework, facilitating its erosion and re-sedimentation (F8a, b). Outer platform. This facies association groups facies F11 and F12 (Table 1). Facies F12 is mainly composed of marls and calcareous nodules embedded in marls. The green calcareous nodules in marls are rich in orbitolinids (F10c and F12a, b) with intercalations of bioclastic limestones that can be cross-bedded (hummocky cross stratification, HCS) and locally show sorting (F10b and F11; Fig. 5h). The marls are structureless and contain a diverse fauna including ammonites, bivalves, gastropods, echinoderms (mainly regular echinoids), foraminifers and oysters. Locally, intensive bioturbation is preserved in these facies, especially in F12, (Thalassinoides, Ophiomorpha, Planolites).

(2)

(3)

(4)

Sequence stratigraphy Time control Biostratigraphy. Schroeder (1964, 1979) defined a correlation of the orbitolinid ranges between the Urgonian deposits of the Basque– Cantabrian region (Weisser 1959; Moullade 1960; Bassoulet & Moullade 1962) and the Villarroya de los Pinares section in the Galve sub-basin (Fig. 3). According to this work, the Barremian–Aptian boundary is in the Palorbitolina lenticularis and Rectodictyoconus giganteus zones (respectively the first and the second zone, Fig. 3), located within the first calcareous bed of the Chert Formation, and the lower Aptian to upper Aptian transition occurs within the Orbitolina lotzei zone associated with the first occurrences of Mesorbitolina sp. (zone 5; Fig. 3). In this work orbitolinids have been studied in two additional sections, the Miravete and Aliaga section (Fig. 3). This information and the literature provide the following biostratigraphic age constraints: (1) At Villarroya de los Pinares (Fig. 3), the occurrence of Palorbitolina lenticularis and Rectodictyoconus giganteus at the base of the calcareous Chert Formation are retained as

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the Barremian/Aptian boundary. The occurrence of Orbitolinopsis praesimplex within the upper part of this formation indicates an early Aptian age. The orbitolinids distribution within the Chert Formation at Miravete and Aliaga sections (Fig. 3) confirms the ages determined at Villarroya de los Pinares. Ammonites (Deshayesites deshayesi and Dufreynoyia furcata) and orbitolinids [Praeorbitolina Cormyi and Palorbitolina lenticularis (advanced species)] found in the Miravete and Villarroya de los Pinares sections indicate an early Aptian age for the Forcall Formation. Occurrences of Mesorbitolina sp., indicative of the upper Aptian, within the Villarroya de los Pinares Formation suggest that the lower Aptian to upper Aptian boundary is located at the base of this formation in Villarroya de los Pinares section. An uppermost Aptian age is attributed to the Benasal Formation based on the presence of a specimen of Acanthoplites bergeroni (Schroeder 1964) within its lower part in the Villarroya de los Pinares section (Fig. 3). The position of the Aptian –Albian boundary can only be bracketed by this ammonite and the typical Albian spore, pollen and benthic foraminifer species found within the overlaying Utrillas Formation (Canerot 1974; Canerot et al. 1982).

Chemostratigraphy. The carbon isotope curve obtained in the Miravete outcrop section is calibrated against the biostratigraphy of benthic foraminifers (orbitolinids) and ammonites presented above. The curve, which covers most of the Aptian time interval, can be subdivided into seven segments (Fig. 7): (1) The first two samples, measured in siliciclastic carbonate channels within the Artoles Formation, are characterized by low d18O values respectively –6.86‰ and –6.05‰ and d13C fluctuating around 1‰ + 0.5‰. An abrupt drop of d13C values and a low d18O value coincide with the boundary between the Artoles and Chert Formations. (2) Low d13C values (d13C ¼ þ1‰ to þ2.7‰) in segment 2, at the base of the Chert Formation

Fig. 5. Aptian microfacies of Galve sub-basin; scale bar is 500 mm: (a) Mixed carbonate-siliciclastic sandstone with intraclast and bioclast debris (facies F3b), Chert Formation, Aliaga section (F3b); (b) Miliolid packstone with some green algae and foraminifers, Chert Formation, Aliaga section (F4); (c) Bioeroded rudist floatstone with sponges, Villarroya de los Pinares Formation, Villarroya de los Pinares section (F6b); (d) Oolitic Grainstone, Chert Formation, Aliaga section (F7a); (e) Oolitic iron-rich packstone with bivalve debris, base of Chert Formation, Aliaga section (F7a); (f) Orbitolinid packstone-grainstone, Chert Formation, Miravete section (F10a); (g) Orbitolinid packstone, base of Villarroya de los Pinares Formation, Villarroya de los Pinares section (F10b); (h) Echinoderm, orbitolinid and bivalve packstone, Forcall Formation, Aliaga section (F11).

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Fig. 6. (a) Photograph of the Miravete section showing lithological formations and medium scale sequences: (b) dominant microbialite encrusting association, mainly composed of Bacinella and Thaumatoporella, scale bar is 500 mm; (c) coral-stromatoporoid-microbialite reef facies at the base of the Villarroya de los Pinares Formation: showing an initial phase of vertical accretion with bioherms (R.) (F9) and inter-biohermal carbonate sands (F8) and a second stage where these buildups spread out laterally to form biostromes several kilometres in size (F9); (d) muddy rudist-dominated floatstone (F6a); Villarroya de los Pinares Formation, scale bar is 5 cm; (e) microbial encrustation (Th.; Thaumatoporella) over a solitary coral fragment (Facies F9b), Villarroya de los Pinares Formation, scale bar is 250 mm; (f) microbialites (Lithocodium, Bacinella) encrusting a coral structure (F9c), Base of Villarroya de los Pinares Formation, scale bar is 250 mm; (g) microbialite (m.) encrusting a stromatoporoid structure (facies F9c); Villarroya de los Pinares Formation and frequently observed in Aliaga section, scale bar is 250 mm; (h) distribution of microbial morphology along a proximal to distal transect in a coral-microbial rimmed platform (see above for illustrations: 2: Fig. 6b; 6: Fig. 6f; and 7: Fig. 6g).

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(3)

(4)

(5)

(6)

(7)

near the Barremian–Aptian boundary, progressively evolve in the lower Aptian to higher values remaining invariant around þ3‰ +0.5‰. This interval is marked by slightly less positive d13C values (d13C ¼ þ1‰ to þ2.8‰) with no significant variation of d18O values. Ammonites belonging to the D. deshayesi zone (Bedoulian) are found in this interval (Fig. 3). A significant positive carbon isotope excursion (d13C values . þ4‰) is observed in the basal part of Villarroya de los Pinares Formation. The occurrence of T. Bowerbanki in the upper part of the Forcall Formation at Villarroya de los Pinares section, suggests an early Aptian age. This segment shows a step-wise decrease of d13C values (d13C ¼ þ1.5‰ to þ3.8‰) with a high variability, which is coupled with a significant drop and rapid fluctuations of d18O values. Segment six is marked by a second major positive carbon isotope excursion with d13C values mostly fluctuating around þ4‰ and þ5‰. Laterally, this interval is correlated with the occurrences of Mesorbitolina sp. in Villarroya de los Pinares section, suggesting a late Aptian age. In segment 7, towards the base of Benasal Formation, these values are progressively replaced by less positive values peaking on a plateau where they oscillate around þ3‰ and þ4‰ and then drop below þ3‰. 13

Although most d C values fluctuate around the Aptian carbonate marine values (þ2‰ to þ3‰) (Fig. 8), there are two stratigraphic intervals that have clear petrographic evidence of diagenetic alteration: the top of segment 1, at the base of the Chert Formation (Fig. 8), which is marked by a strong negative shift in d13C, and segment 5, which shows a trend towards lighter values. These will be discussed further below. Sequence stratigraphic model. The sequence stratigraphic model is based on the correlation of five outcrop sections, two of which are presented in detail in Figures 7 and 9. Four types of characteristic stratigraphic surfaces are distinguished: (1) flooding surfaces, indicated by the presence of glauconite and a net increase in the content of clay and pelagic fossils above the surface; (2) hardgrounds which display intense burrowing and boring; (3) sub-aerial exposure surfaces, indicated by meteoric and vadose cements; and (4) erosive transgressive surfaces above which there are ferruginous ooids, glauconite and phosphates-rich deposits (Figs 9 &

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10). In addition, evidence for tectonic control on depositional geometries and syn-sedimentary faulting were provided by field photo panels and aerial photographs (Vennin & Aurell 2001). A subdivision at four scales of sequences is made (Fig. 10): (a) at the large scale, a major marine episode (Chert, Forcall, Villarroya de los Pinares and lower part of Benasal Formations) between two episodes of continental/estuarine sedimentation (Artoles Formation and upper part of Benasal Formation) is recorded with a maximum flooding within the Forcall marly Formation (MFS1) (Fig. 10). This sequence is bounded by two major erosive surfaces SB I-1 and SB III-1, interpreted as periods when accommodation was at a minimum. (b) The presence of a distinct sub-aerial exposure at the lower/upper Aptian boundary, in the middle of the marine succession, allows subdivision of this overall trend into two large scale, second-order, sequences (I and II; Fig. 10): the first one shows an overall deepening-up trend from a siliciclastic continental environment to a basinal marine environment, followed by shallowing-up on a carbonaterimmed platform. The second one shows a deepening trend from mud flat to an outer ramp setting, followed by a shallowing-up into terrestrial siliciclastic deposits. (c) These sequences can be further subdivided into third-order sequences, referred to as sequences: I-1, I-2 and II-1, II-2 which themselves are composed of (d) fourth-order sequences (Figs 7 & 10). The evolution of the third-order sequences is described in detail below. Sequence I-1. Sequence I-1 consists of a composite deepening–shallowing up cycle. It comprises mainly siliciclastic sediments in the lower part, and is carbonate-rich in the upper part, dominated by orbitolinids and ooids (Fig. 10). The fluviatile, stacked channels at the base of the Artoles Formation represent the base of the first sequence (I-1). The channels have erosive bases (SB I-1) and are of mixed carbonate-siliciclastic composition (Table 1). In vertical continuity, the deposition of the Chert Formation especially in SE sector (Miravete, Portoles, and Villarroya de los Pinares sections) records an evolution to more pronounced marine conditions indicated by a main transgressive surface (TS I-1). The top of this sequence (SB I-2) is bounded by an exposure surface in the proximal area and a bioturbated horizon (Ophiomorpha, Thalassinoı¨des and Planolites) and/or a hardground surface in the distal area of the platform. Updip, the sub-aerial exposure surface shows fenestrae structures and meteoric to vadose cements (F4). The hardground surface encrusted by corals, serpulids and algae (F7c) and the bioturbated horizon are interpreted as a temporary halt in sedimentation.

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Fig. 7.

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Fig. 8. Cross plot of d18O v. d13C for the 106 bulk-rock samples collected in Miravete. The facies interpretations are indicated for each sample and refer to Table 1. Most values plot around presumed marine values (d13C ¼ þ2‰ + 3‰ and d18O ¼ 23‰). One sample located near the transgressive surface TS in S I-1 (a circle in the diagram) shows a strong negative shift in d18O and d13C due to high content of early diagenetic products. Twelve samples (area delimited by a dashed line) also show a significant negative shift in d18O. These samples are located below a major sub-aerial exposure surface (SB II-1) and marked by petrographic evidence of diagenetic overprint. The carbon-isotopic curve in this ‘anomalous’ interval also reveals a significant decrease of the d13C values and a rapid fluctuation of both carbonand oxygen-isotope records (segment 5 in Fig. 11).

The early transgressive systems tract (TST) is characterized by a continental to coastal plain depositional environment. These deposits are composed of continental lacustrine marls and limestones rich in ostracods and charophytes (F2), pedogenetized horizons (F1) and mixed siliciclasticcarbonate channels indicating fluviatile settings (F3a). The increasing number of bioclastic limestone beds rich in bivalves and gastropods (F5a, b and c) towards the top of this interval points to a progressive evolution from a coastal plain environment to a more marine-dominated environment. The early

TST deposits are separated from late TST deposits by a transgressive surface (TS I-1, Fig. 10). This surface is marked by the occurrence of glauconite and iron oxides and is overlain by ferruginous ooids (Aliaga and Villarroya de los Pinares sections) (Fig. 5e). An alternative sequence stratigraphic interpretation is to place sequence boundary I-1 at this surface. Pending further dating of the underlying coastal plain deposits, we follow here the former interpretation. The late TST deposits are characterized by carbonate sedimentation (F7a, b and c –F10a and b)

Fig. 7. Miravete outcrop section illustrating Aptian facies evolution, oxygen- and carbon-isotope stratigraphy in a platform setting. The medium and small scale sequences are shown (Location in Fig. 1; legend in Fig. 4; the colour of the facies are in Table 1). Sequence S I-1 shows orbitolinid bioaccumulations in the transgressive trend and oolitic-bioclastic shoals in the regressive trend. Sequence S I-2 consists of orbitolinid and ammonite-rich marls in the transgressive trend, and coral-stromatoporoid-microbialite in both regressive and forced regressive trends. Sequence S II-1 mainly corresponds to a rudist-rich mud flat in both transgressive and regressive trends. Sequence S II-2 deposits are mainly composed of limestone rich in orbitolinids in ramp setting. The increasing amount of channelized mixed siliciclastic-carbonate sandstones towards the top of the regressive trend indicates a period of high detrital input in the Galve sub-basin.

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Fig. 9. Aliaga section: outcrop photographs and palaeoenvironmental interpretations based on macrofacies, microfacies and bedding pattern analysis. The colour code for the facies is in Table 1. Lithostratigraphic nomenclature and sequence subdivisions are shown (Location in Fig. 1).

Fig. 10. Facies and sequence correlation scheme of the five studied outcrop sections the Galve sub-basin. The upper boundary of sequence S II-1 is taken as reference level. The distribution of facies and depositional environments are indicated and refer to legends in Table 1 and Figure 4.

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in the SE sector (Miravete, Portoles, and Villarroya de los Pinares) and mixed siliciclastic-carbonate deposits (F3b–F7a to c– F5b and c) towards the NW sector (Aliaga and Peral sections; Fig. 10). During the last stage of the transgression, slightly deeper facies with evidence of storm reworking were deposited over the platform and have been interpreted as open marine facies in the shoreface to upper offshore environments (mid to outer platform settings): orbitolinid-rich wackestones (F10b) in the Aliaga, Peral and Miravete sections and orbitolinid–rich marls (F10c) in the Portoles and Villarroya de los Pinares sections. The top of the general deepening-upward trend recorded in the TST is marked by a condensed surface interpreted as a maximum flooding surface (MFS I-1). This condensed surface is strongly bioturbated with various ichnological associations: (1) Thalassinoı¨des, Planolites in the facies F5c and F10a and (2) Teichichnus, Thalassinoı¨des and Phycodes in the facies F10c. In the distal SE area, the subsequent shallowingupward trend corresponding to a highstand system tract (HST) is marked by the evolution from orbitolinid-dominated limestones and marls (F10b and c) to oolitic-bioclastic shoal deposits (F7) towards the top of the sequence (SB I-2; Fig. 10). In the more proximal Aliaga and Peral sections, the mixed siliciclastic-carbonate sandstones (F3b) and the bioclastic packstones and floatstones (F7b and d–F5b and c) are overlain by packstones and wackestones rich in miliolids, bivalves and rare ooids (F4b) interpreted as protected lagoonal facies deposited behind the ooid-bioclastic shoals. In Aliaga and Peral, the mixed siliciclasticcarbonate sandstones (F3b) show abundant obliquestratification to rarer, plane or low angle stratifications in the bioclastic carbonate layers (F7b, d). The cross-bedding stratification and fining-upwards sequences observed within the sandstone deposits indicate high energy conditions. The orientation of foresets suggests a siliciclastic supply from the west. Within this sequence, a stacking pattern of metre-scale discontinuous genetic units (Fig. 10) is observed and corresponds to the overall migration of the bioclastic sand shoals (Fig. 4a). The origin of these units is related to the hydrodynamic processes which affected the carbonate ramp. Within the transgressive system tract, the genetic units onlap on the transgressive surface (TS I-1) resulting in the landward shift of the palaeoshoreline. The genetic units belonging to the HST are bounded by erosional surfaces and composed of partially eroded layers prograding towards the SE. The hummocky cross-stratification and the wave ripple marks observed within this wedge underline the permanent swell and storm influence. In the distal area, the presence of syn-sedimentary normal

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faults (ENE –WSW trending faults) sealed by Sequence I-2 probably contributed to the increase of accommodation space and to the separation of the various genetic units in Sequences I-1 and I-2. Sequence I-2. The sedimentation within this sequence is characterized by carbonate and marly deposits corresponding to the Forcall Formation and to the lower part of the Villarroya de los Pinares Formation (Fig. 10). The sequence (S I-2) exhibits a composite deepening–shallowing cycle. Carbonate production is dominated by orbitolinids, microbialites, corals and rare rudists. The basal sequence boundary (SB I-2) is described above. The top sequence boundary (SB II-1) is a major sub-aerial exposure surface in the proximal domain, associated with vadose and meteoric cements. Downdip this correlates to a sharp facies change from orbitolinid-dominated green marls (F12a) to bioclastic grainstones (F7d) in Portoles and orbitolinid-dominated wackestones in Villarroya de los Pinares (F10a). This surface marks a significant change in the sedimentary facies and is interpreted as a basinward shift of the sedimentary system. Sequence I-2 is made up of three systems tracts: a TST, an HST and a forced regressive systems tract (FRST) (Hunt & Tucker 1995). Sequence I-2 is composed of 10 highfrequency sequences, metre–decimetre thick, correlatable along the Galve sub-basin (Fig. 10). The early TST is characterized by upper shoreface to offshore facies (inner, medium and outer platform depositional environment). The lower part of the TST consists of platform interior mudstones and wackestones rich in miliolids (F4a) (upper shoreface) passing downdip in Villarroya de los Pinares (Fig. 10) to bioclastic packstones (F5c) and orbitolinid-dominated packstones to wackestones (F10b) interpreted as mid-platform environments (lower shoreface) with grainy tempestite deposits (F5c). Direct evidence of onlap of highfrequency sequences on the SB 1-2 surface is observed between the Peral and Miravete sections (Fig. 10). During late transgression, orbitolinid-rich facies (F10a to c) and green marls (F12a and b) dominate, which have been interpreted as mid to outer platform environments (lower shoreface to upper offshore). These deposits are locally replaced by coral and microbial patches (F9a). In Peral, a significant lateral change of thickness in the orbitolinid-dominated marls and limestones (F10b and c) has been related to the activity of the synsedimentary ENE –WSW-trending normal faults affecting both Sequence I-1 and I-2 (Vennin & Aurell 2001). The maximum flooding surface (MFS I-2) is marked in the Miravete to Villarroya de los Pinares sector by a condensed surface, overlain by

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ammonite and brachiopod rich-marls (F12b –F11). In Aliaga and Peral, this surface passes updip to an intensely bioturbated horizon (ichnological association with Thalassinoı¨des and Planolites). The maximum flooding period recorded in the Forcall Formation corresponds to the evolution to a dominantly marly sedimentation in the Galve subbasin and to a period of maximum flooding described at a regional scale in the Maestrat basin (Vennin et al. 1993; Soria et al. 1994; Salas et al. 1995; Vennin & Aurell 2001). The HST deposits consist mainly of bioturbated green marls and limestones, rich in orbitolinids (F10a to c –F12a) in the distal platform (Fig. 3b). In the proximal Aliaga and Peral sections, these orbitolinid-rich deposits are locally replaced by intercalations of coral-microbialite reef facies (F9c) prograding basinward (Fig. 10). At this time a clear platform to basin topography was established with a boundstone-rimmed margin and the facies association and the depositional geometries observed in this sequence correspond to depositional model B (Fig. 4b). The top of the HST marks the base of the forced regressive deposits, which are well exposed in Aliaga, Peral and Miravete sections (Fig. 10). This wedge is characterized by a significant offlap migration of stromatoporoid-coral and rare microbialite reef facies basinward and corresponds to an abrupt regional decrease of the relative sea-level in the Maestrat Basin (Canerot 1974; Salas & Guimera 1996). Along the Miravete section, the FRST deposits comprise coral-stromatoporoidmicrobialite reef facies (F9b and c– F8) overlain by miliolid- and rudist-dominated wackestones (F4a and b–F6a). The reef facies first aggraded (bioherm growth phase) and then prograded (biostrome growth stage) basinward (Fig. 3b). In the proximal Aliaga and Peral sections, the FRST deposits consist mainly of wackestones to packstones rich in miliolids and rudists which have been interpreted as representing a protected lagoonal environment (F6a) (upper shoreface) deposited behind the reef. Towards the SE (Portoles and Villarroya de los Pinares), the reef facies are laterally correlated with deeper-water orbitolinid-rich marls (F12a). Sequence II-1. Sequence II-1 consists of a deepening–shallowing cycle. Carbonate production is dominated by rudists, gastropods and miliolids. Sequence II-2 corresponds to the upper part of Villarroya de los Pinares Formation (Fig. 10) and is composed of six high-frequency sequences, metre– decimetre thick, correlatable along the Galve subbasin (Figs 7, 9 & 10). The sequence is bounded at its base by sequence boundary SB II-1 described above. The top of this sequence is a hardground surface (SB II-2) overlain by an iron oxide crust

and penetrated locally by borings. This sequence consists of a lowstand system tract (LST), a TST, and an HST. The platform to basin topography developed during the previous sequence I-2 is onlapped by LST deposits (Fig. 10). This lowstand wedge can be traced downslope from Portoles to Villarroya de los Pinares section and is composed of bioclastic grainstones (F7d) passing downdip to orbitoliniddominated wackestones to packstones (F10a and b) rich in bivalves and rudist debris (Fig. 4c). The top of the LST is a bored hardground and is interpreted as a transgressive surface (TS) (Fig. 10). The early TST deposits overlying the transgressive surface are characterized mainly by muddy wackestones and packstones rich in Chondrodonts bivalves and rudists (F6a and b) interpreted as indicating a shallow platform top environment (upper shoreface). The rudists are represented by Requienids (Toucasia sp.) and Capronitids (Monopleura sp. and Horiopleura sp.). In the NW Aliaga section, the facies are marked by a significant enrichment in miliolids and micritic matrix (F4b). The local hummocky cross-stratification observed in the most proximal part of the platform indicates the episodic influence of storms. The common occurrence of orbitolinids-rich marls (F10c) within the platform rudist –rich facies highlights the occasional communication with more open marine depositional environments. The transgressive deposits in the vicinity of the Portoles section are thicker compared to the rest of the platform and only composed of green marls rich in orbitolinids (F10c). This lateral thickness and facies variation over a distance of 20 –50 m is interpreted as a local palaeotopographic low inherited from the displacement of syn-sedimentary faults observed in this area. The top of the TST consists of orbitolinid-rich marls and wackestones (F10c) interpreted as a maximum flooding period. The overlying HST deposits are mainly facies dominated by rudists, miliolids and gastropods (F4a –F6b). The thickness of the rudist-rich limestones tends to increase downdip towards Peral and Villarroya de los Pinares sections. The diversity of the rudists is strongly reduced and the rudist-rich accumulations are mainly composed of Requienidae (Toucasia sp.). The facies association and the depositional geometries observed in this sequence correspond to Depositional Model C, the rudist-dominated mud-flat (Fig. 4). Sequence II-2. This sequence consists of a deepening–shallowing cycle, and is dominated by carbonate in the lower part, whereas in the upper part there is a gradual increase of siliciclastic sediments. This sequence corresponds to the Benasal Formation.

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The sequence is bounded at the base by a hardground surface (SB II-2), described above, and at the top by an erosional surface (SB5) overlain by a thick channelized sand system, rich in glauconite (Fig. 10). In Aliaga to Miravete area, the TST deposits consist of wackestones rich in benthic foraminifers, green algae, serpulids and rudist debris (F4b) intercalated with orbitolinid-dominated packstones (F10a). This facies association, interpreted as upper shoreface depositional environment (inner ramp), passes downdip to dominantly marly deposits rich in orbitolinids (F10c) (lower shoreface to upper offshore facies, mid to outer ramp depositional environment) with intercalated decimetre-thick colonial coral patches (F9a) (Fig. 4a). The top of the TST is marked by a maximum flooding surface/ period (MFS II-2) corresponding to a bioturbated horizon encrusted by oysters in Aliaga and downdip to the deposition of orbitolinid-rich marls in the Miravete to Villarroya de los Pinares area (Fig. 10). The overlying HST is marked by high-energy bioclastic shoal deposits (F7a and d) and orbitolinidrich packstones (F10a). The increasing amount of channelized sandstone deposits towards the top of the HST reflects a period of enhanced siliciclastic input in the Galve sub-basin. The facies association in this sequence corresponds to Depositional Model A, the mixed ramp (Fig. 4a). Despite no direct evidence of onlap geometries or major facies change observed in the field, the detailed mapping of this sequence on aerial photographs reveals some local subtle changes in thicknesses (Vennin & Aurell 2001) which could be related to syn-sedimentary extensional structures.

Discussion Chemostratigraphic correlations The d13C curve obtained in the Galve sub-basin is compared to other curves in pelagic and benthic settings along the NeoTethys in Figure 11. The following observations are made. Segment 1 and 2 show the negative peak followed by an overall increasing d13C trend which is typical for the lower part of the lower Aptian, and has been observed in many shallow water and pelagic domains (Fig. 11). This minor C isotope event follows the eustatic sea-level change, with increasing d13C values associated with the early Aptian transgression (Jarvis et al. 2002). The slightly decreasing d13C trend observed in segment 3 (Fig. 11) occurs in the D. deshayesi ammonite zone, and has been previously recorded in high resolution carbon isotope records in the NeoTethys, where it coincides with the occurrence of widespread organic carbon-rich horizons, Oceanic Anoxic Event 1a, in the pelagic domains (‘Livello

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Selli’ in the southern Alps of northern Italy, Coccioni et al. 1987, and ‘Niveau Goguel’ in the Vocontian basin, Weissert & Breheret 1991). The origin of this event is subject to debate. Some authors suggest a variation of carbon cycling (Menegatti et al. 1998), volcanic activity (increasing volcanic Ontong Java and Manihiki Plateau activity) and hydrothermal activity (Bralower et al. 1994, 1999; Leckie et al. 2002; Ingle & Coffin 2004) or thermal destabilization of methane hydrate (Jahren et al. 2001; Beerling et al. 2002). The transition to a major positive shift in d13C, in segment 4 (Fig. 11), is dated as late early Aptian (D. deshayesi and D. Furcata ammonite zones). This positive excursion, well-documented in shallowwater carbonate platform and pelagic settings, occurs during a global sea-level rise, and is interpreted as a consequence of an accelerated rate of Corg and Ccarb burial (Weissert 1989; Weissert et al. 1998) and is possibly a result of slight variations in the oxygenation of the bottom waters from anoxic to suboxic (Mutterlose et al. 2009). Segment 5 is more difficult to interpret. On the one hand, the trend to lower d13C values, dated as the early/late Aptian transition by the Orbitolinids may correspond to a documented excursion to less positive d13C values within the G. ferreolensis and G. algerianus foraminifer zones. This event is interpreted as a new stable mode of carbon cycling with increasing carbonate sedimentation rates, a decrease of erosion rate and nutrient supply in oceanic domains and/or a cooling event (Weissert & Lini 1991; Fo¨llmi et al. 1994; Jenkyns 1995; Weissert et al. 1998; Bralower et al. 1999; Wortmann et al. 2004). On the other hand, this interval is located below a major sub-aerial exposure surface and the lower d13C values may result from an early diagenetic influence. Although the covariance of d13C and d18O composition of the carbonate rocks in this interval is low (– 0.034), the isotopic curves show a significant decrease of the d13C values accompanied by a decreasing oxygen isotope trend and a rapid fluctuation of both isotopic compositions. The diagenetic impact on the isotope measurements is supported by evidence of meteoritic cements in some of the corresponding thin sections. Nevertheless, as proposed by Gro¨tsch et al. (1998) in such a case, the least negative d13C values should correspond to the less affected original marine isotope signal. If we accept this assumption, the most positive d13C values (around þ2‰ to þ3‰) could have preserved the remaining original marine signal. The identification of a sub-aerial exposure towards the top of this interval also suggests a gap in the sedimentary succession and thus a discontinuity in the carbon and oxygen isotope record. The positive trend observed in segment 6 may correlate with the late Aptian positive excursion

132 J.-C. EMBRY ET AL. Fig. 11. Correlation of the Miravete outcrop section with the sections published by van Buchem et al. (2002), Menegatti et al. (1998), Moullade et al. (1998), Scholle & Arthur (1980), Masse et al. (1990), Embry (2005), Herrle et al. (2004), Bralower et al. (1999) and Jenkyns et al. (1995) from different depositional environments. All published sections presented in this figure are calibrated against biostratigraphy of planktic foraminifera, calcareous nannofossils and/or ammonites (Gradstein et al. 2004). For reference the oceanic anoxic events (1a and 1b), the Niveau Jacob (NJ) observed in the Vocontian Basin (Weissert & Breheret 1991), volcanic activity and sea-level changes (Haq et al. 1988) are presented to highlight the major global environmental changes occurring during the Aptian.

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recognized in other localities and occurring in the upper G. algerianus and H. trocoidea foraminifer zones (d13C . þ4‰) (Fig. 11). This event occurs after the early/late Aptian termination event of many Aptian platforms (Fo¨llmi et al. 1994; Gro¨tsch et al. 1998; Herrle et al. 2004; Jenkyns 1995; Masse & Philip 1981; Vahrenkamp 1996; Weissert & Lini 1991). Within this interval, the isotopic curves are locally disturbed by negative shifts of the d13C values and positive or negative deflection of the d18O signal. In segment 7, a gradual decrease of d13C values is observed (Fig. 11) which coincides with the d13C trends published from upper Aptian pelagic successions in the Pacific and Tethyan domains (Weissert & Lini 1991; Bralower et al. 1999; Jenkyns 1995; Herrle et al. 2004). To summarize, the carbon isotope record of the Miravete section fits well with the published carbon curves for the benthic and pelagic domains. It provides a rare example of a preserved upper Aptian shallow-water record.

Relative sea-level change and depositional sequences The Aptian sequences and relative sea-level curve proposed for the Galve sub-basin (Fig. 12) are compared to three other Aptian shallow-water platform domains along the northern and southern Tethyan margin (Fig. 13): the Provence area (Masse et al. 1999); the Vercors platform interior domain (Hunt & Tucker 1993; Arnaud et al. 1998) and the Oman and Abu Dhabi platforms (van Buchem et al. 2002, 2009). The correlations between these sites are made using the currently available, but limited, biostratigraphic and chemostratigraphic data, in the shallow-water carbonate environments. The following observations are made. SB I-1 (latest Barremian/earliest Aptian). In the four compared areas a sequence boundary is observed in the upper part of the upper Barremian, or possibly the lowest part of the lower Aptian. Biostratigraphic resolution is not sufficient to date these surfaces precisely, but in all locations they are marked as a clear change in sedimentation pattern. In the Galve sub-basin, SB I-1 is tentatively placed at the base of a stacked channel succession in a continental and coastal plain setting, that shows a deepening upward trend above (Fig. 10). In the Provence, a sharp change from shallow-water carbonate platform top sediments to intrashelf basinal facies is observed in the latest Barremian (ammonite dating; Masse 1993). In the Vercors, this sequence boundary is marked by a sharp facies change, from a grainy, rudist-dominated platform top facies, to channel incision and fill with argillaceous,

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orbitolinid-dominated facies (lower Orbitolina Beds; Arnaud et al. 1998). In Oman, a comparable facies change is observed, from the rudistdominated upper Kharaib platform deposits, to the argillaceous, orbitolinid-dominated Hawar facies (van Buchem et al. 2002). The thick succession of fluvial coastal plain sediments observed in the Galve basin, may reflect either a local high subsidence rate during the late Barremian to early Aptian, possibly related to the active rifting at that time (Liesa et al. 2006), or the difficulty of precise interbasinal correlation of sequence boundaries due to limitations of the available time control. SB I-2 (early Aptian). Sequence boundary I-2 also has a different expression in the different localities. In the Galve basin, it corresponds to a surface on top of a shallowing-upward trend in a mixed carbonate-siliciclastic ramp setting, with evidence for exposure in the proximal domain (Fig. 10). In Provence (Masse et al. 1993; Embry 2005) and the Vercors (Arnaud et al. 1998), at approximately the same time, carbonate production ceased. This was accompanied in the Vercors by channel incision. A similar demise of the shallow-water carbonate platforms is reported from Switzerland (Fo¨llmi et al. 1994, 2006; Weissert et al. 1998; Embry 2005). In Oman, at this time, a subtle change in sedimentation is observed, mostly determined by the increased sea-level rise that followed the deposition of the orbitolinid-rich Hawar facies (van Buchem et al. 2002, 2009). MFS I-2 (late early Aptian). The best dated event is the late early Aptian eustatic sea-level rise which coincided with OAE 1a deposition in the pelagic realm (e.g. Fo¨llmi et al. 1994, 2006; Weissert et al. 1998). A remarkable regional difference in sedimentological expression is observed at this time. This sea-level rise caused the initiation of intrashelf basins on the Arabian Plate (e.g. Murris 1980; van Buchem et al. 2009), and the timeequivalent aggradation of microbial-dominated facies in the shallow-water domains (Pittet et al. 2002; van Buchem et al. 2002; Hillgartner et al. 2003; Immenhauser et al. 2005). In the Galve basin, a general deepening to marly, offshore facies was observed, whereas in the Provence basin, bioclastic-glauconitic facies were deposited at this time (Masse 1993). In the Vercors basin, only siliciclastic sediments and orbitolinid-filled incised channels (upper Orbitolina Beds) are recorded (Arnaud et al. 1998), as in the Helvetic Alps (Follmi et al. 2006). SB II-1 (early –late Aptian boundary). Sequence boundary II-1 occurs at the lower –upper Aptian

134 J.-C. EMBRY ET AL.

Fig. 12. Chronostratigraphy of Barremian –Aptian in Galve sub-basin. The depositional environments are indicated. Time-scale is based on Gradstein et al. (2004). The carbon isotope curve from the Miravete outcrop section is shown. The best biostratigraphic controls are shown and correspond to: (1) occurrences of Palorbitolina lenticularis,

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boundary in the Galve basin, where there is evidence for a forced regression and sub-aerial exposure below this surface. The facies of this platform margin are dominated by corals and microbialites. A detailed study of this forced regressive event suggests that the amplitude of the regional relative sea-level fall in the Galve sub-basin was approximately 60 m (Bover-Arnal et al. 2009). In the Provence basin, this surface lies on top of the bioclastic, glauconitic facies (Masse 1993). On the Arabian Plate, the early –late Aptian boundary marks the onset of sea-level fall and platform exposure, which reaches its climax at the end of the Subnodosocostatum zone, with a sealevel drop of at least 30–40 m, which terminated the early Aptian platforms (van Buchem et al. 2010). This sea-level fall is considered to be of glacio-eustatic origin, caused by a climatic cooling in the late Aptian (Puce´at et al. 2003; Steuber et al. 2005). Whereas sedimentation on the Arabian plate remained restricted to the intrashelf basins (van Buchem et al. 2010), carbonate deposition resumed after this exposure event in the Galve basin, with the deposition of a rudist-dominated (Taucasia) platform top succession. These deposits are possibly time-equivalent with the deposition of the pelagic Gargas marls in the Provence basin (Masse 1993), and glauconitic grainstones of the Lumachelle Formation in the Vercors basin (Hunt & Tucker 1993), but further biostratigraphic work is needed to substantiate this. SB II-2 late Aptian. In Spain, this sequence boundary marks the end of the rudist platform, which is subsequently replaced by mixed carbonatesiliciclastic, orbitolinid-dominated deposits. Based on the carbon-isotope curve this surface may be placed in melchioris/nolani zones. In Oman, at the same time, low-angle prograding clinoforms were deposited within the intrashelf basin (Maurer et al. 2010; Pierson et al. 2010). The uppermost Aptian time equivalent deposits have not been preserved in the Provence and Vercors platform systems due to Cretaceous erosional episodes (upper Aptian and/or Mid- to Upper Cretaceous) and/or nondeposition (Masse et al. 1981; Hunt & Tucker 1993; Arnaud et al. 1998). The local expression of the shallow water, rudistic platform facies in the

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Galve basin may be a result of tectonic subsidence, creating the necessary accommodation. SB III-1 latest Aptian. Sequence boundary III-1 is in Spain, an erosional surface at the base of a glauconitic, siliciclastic unit, with bidirectional crossbedding and oysters banks. These tidally- influenced coastal plain deposits, mark a return of the siliciclastic system. No clear link with the succession recorded in Oman is evident at this time. To summarize, the following general trends have been observed: (1) The upper Barremian carbonate platforms (‘Urgonian facies’) were terminated in the earliest Aptian along the northern and southern NeoTethys margins. This was a virtually isochronous event, controlled by a eustatic drop in sea-level. (2) During the following early (slow) transgression, orbitolinids were deposited in argillaceous to muddy facies in all studied locations. In the Provence and the Vercors successions, this facies was quickly replaced by a rudistdominated shallow-water carbonate facies, whereas in Oman and Spain the orbitoliniddominated facies continued. (3) The worldwide-recognized early Aptian transgression (OAE 1a) is well expressed on the Arabian Plate and in Spain, where a retrogradational and aggradational sedimentation pattern are observed. In the Provence and Vercors systems the termination of carbonate sedimentation is observed just before or at this time. A second phase of platform termination occurred in the early late Aptian, when the platforms on the Arabian Plate became exposed. At the same time, or somewhat later, sedimentation on the last rudistdominated platforms in Spain also ceased. The causes for this stepwise, diachronous, termination of Aptian platforms are the subject of debate (Weissert et al. 1998; Skelton & Masse 2000; Skelton et al. 2003). (4) Microbial facies dominated at different times: on the Arabian plate they influenced sedimentation on the platform during the early Aptian transgression (van Buchem et al. 2002; Immenhauser et al. 2005) and remained an

Fig. 12. (Continued) Rectodictyoconus giganteus and Orbitolinopsis praesimplex indicating an early Aptian age within the Chert Formation in Villarroya de los Pinares and Aliaga sections; (2 and 3) ammonites of early Aptian age (Deshayesites deshayesi and Dufreynoyia furcata) found within the Forcall Formation in the Miravete and Villarroya de los Pinares sections; (4) occurrences of Mesorbitolina sp., indicative of the upper Aptian, at the base of the Villarroya de los Pinares Formation in Villarroya de los Pinares section; and (5) presence of Acanthoplites bergeroni of latest Aptian age (Schroeder; 1964) in the basal part of the Benasal Formation in Villarroya de los Pinares section. The chronostratigraphic scheme highlights diachroneity and the gaps of deposits near the sequence boundaries and the transgressive surfaces.

136 J.-C. EMBRY ET AL. Fig. 13. Summary chart of the sequence-stratigraphic interpretation of the Aptian deposits of this study and comparison with other regional/global trends. (a) Time-scale from Gradstein et al. (2004). (b) Simplified chronostatigraphic schemes with facies distribution and platform architecture through time for the Galve sub-basin, the Provence platform (Masse et al. 1993), the Vercors platform interior domain (Hunt et al. 1993; Arnaud et al. 1998) and Oman platform (van Buchem et al. 2002, 2009). (c) Volcanic activity (Larson 1991) and stratigraphic position of the oceanic anoxic events 1a and 1b (Weissert et al. 1998). (d) Eustatic sea-level and T –R cycles from Haq et al. (1988) re-calibrated by Gradstein et al. (2004).

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(5)

important component along the southern Arabian Plate margin (Hillgartner et al. 2003). They were a dominant component in the late early Aptian platform margin in the Galve basin, but only a minor component in the early Aptian carbonate platforms of the Provence and Vercors systems. The thick Aptian succession in the Galve basin (approximately 380 m), as compared to the much thinner successions in Provence (maximum 250 m), Vercors (around 80 m in the platform interior settings) and Oman (110 m), suggests that tectonic subsidence created much more accommodation space in this location, thereby allowing for a better preservation and expression of the Aptian sequences. The measured carbon and oxygen isotope curves can, therefore be considered as a good reference for the Aptian signature.

Fossil community response to global and local controlling factors A striking feature of the Aptian shallow-water carbonates is the occurrence of three dominant faunal associations: the orbitolinid-dominated muddy and argillaceous facies, the microbial/coral boundstone facies, and the muddy to grainy rudist-dominated facies. These are discussed in more detail below. Orbitolinids. Orbitolinid-dominated assemblages occur preferably in mixed siliciclastic-carbonate systems throughout the peri-Tethyan platform domains (Arnaud-Vanneau & Arnaud 1990; Funk et al. 1993; Vilas et al. 1995; Bernaus 2000; Pittet et al. 2002; van Buchem et al. 2002). The ‘orbitolinid events’ that occurred during the early Aptian transgression have been interpreted as a consequence of increased terrigeneous runoff induced by a more humid climate (Weissert 1989; Weissert & Lini 1991) and clay mobilization during flooding of the exposed platform (Pittet et al. 2002). These caused mesotrophic conditions, which were advantageous to the development of the orbitolinids. Orbitolinids have been found in very different environments, such as intertidal deposits, intercalated with miliolid-rich facies and mud-crack features (Oman; van Buchem et al. 2002), (tidal) channel fills (this study and Vercors; Arnaud et al. 1998), and open marine platform margin settings (this study; Oman; Witt & Gokdag 1994; van Buchem et al. 2002). In the Galve basin, a semiquantitative analysis was performed to describe the size, abundance and morphology of the orbitolinids. A relationship between the bathymetry, the nutrient input, the clay proportion in the deposits and orbitolinid development is noticed (Figs 7 & 9). From the early Aptian transgressive surface

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their abundance progressively increases until the maximum flooding of the early Aptian. Then their abundance gradually decreases. Their morphological evolution shows that the largest and more discoidal forms occurred during the maximum flooding period in the Forcall marls. The morphology gradually evolves to more conical and smaller forms during the subsequent relative sea-level fall (Figs 7, 9 & 10). A similar evolution, but less wellexpressed, is also observed within the medium-scale sequences. Orbitolinids increase in Sequence SII-2 (late Aptian) and are frequently observed intercalated with rudist-rich facies and associated with siliciclastic inputs. The symbiosis of these foraminifera with chlorophyllian unicellular algae (Zooxantellae) as purposed by Hottinger (1982, 1997), may explain the evolution toward more discoidal and larger forms in more clay-rich deposits where the maximum water depth was reached. The increased turbidity induced by significant clay input and greater water depth reduced the amount of light in the environment and favoured fast-growing organisms with asexual reproduction (Birkeland 1988), as proposed by Vilas et al. (1995) for large foraminifera such as orbitolinids. The increase of the exchange surface may have been a morphological adaptation to enhance photosynthesis in these light-dependant foraminifera (Hottinger 1997). Their abundance may be related to the nutrient level. Increased nutrient input tends to favour the development of rapidly growing organisms with r-selected traits because they can rapidly respond to a sudden episode of nutrient supply into the environment (Birkeland 1987), such as may occur during periods of transgression. Corals and microbial deposits. In the Galve basin, the dominance of solitary and branching corals associated with Bacinella, Thaumatoporella and Lithocodium microencrusters, is observed in the prograding part of early Aptian Sequence II-2 (Fig. 6c, g). Microbialites show changes in morphologies through the proximal–distal coralmicrobialite-rimmed platform from encrusting forms in Aliaga to rhodolith-like structures in Villarroya de los Pinares (Fig. 6h). In bioherms, microbialites are well developed and coalescent structures are commonly observed (Fig. 6h). This association of builders formed the margin of the platform at that time (Fig. 4). A similar proliferation of microbial facies has been observed in the lower Aptian platforms and notably platform margin of the Shu’aiba Formation on the Arabian Plate (Hillgartner et al. 2003; Immenhauser et al. 2005). In Oman, the climax of the microbial facies deposition was contemporaneous with and just after the OAE 1a crisis, and is suggested to be the

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shallow-water expression of this pelagic event (Hillgartner et al. 2003; Immenhauser et al. 2002, 2005). However, a more specific analysis of the corals and microbialites is needed to assess the trophic level (e.g. Olivier et al. 2008) and therefore the potential control of the nutrient input (as well as water carbonate alkalinity and ‘sea-water saturation with respect to CaCO3’) on the development of these microbialite-rich reef facies occurring just after the OAE1a event (Immenhauser et al. 2005). Sediment influx is an important controlling factor in the distribution of reef organisms and in overall reef development (Hubbard 1986; Macintyre 1988). Reefs are generally better developed, have more coral species, higher coral cover, and faster rates of framework accretion the farther they are from sources of runoff or the lower the sediment load in overlying waters (Morelock et al. 1983; Hubbard 1988). We infer that the observed changes in reef morphologies in the Galve basin from bioherms to biostromes result from relative sea-level fluctuations superposed on the overall decreases in accommodation space during the final progradational of Sequence I-2. External processes (clastic discharge associated with climatic changes and relative sea-level variations) may thus have controlled the variations in progradation rate and carbonate production, with reef productivity increasing with decreasing or lack of runoff. Metricpatches correspond to a very gently sloping platform (Fig. 4b), and are observed either in the back reefs of rimmed-platform (S I-2) or in mud flat depositional environments (S II-1 and S II-2) whose growth is disturbed by clastic input (marls and siltstones). Rudists. The Aptian is the acme for rudist faunas in SE (Vilas et al. 1993) and NE Spain (Masse & Fenerci-Masse 2008), whereas for the Urgonian platforms of SE France this is the Barremian time (Arneau-Vanneau et al. 1982). The monospecific rudist-dominated mud flat environment observed in Sequence II-1 of the Galve basin (Fig. 10) represents the last phase of Aptian rudist-dominated platform sedimentation. The main rudists (Requienid) show low elevation indices (clingers) and achieved high stability by maximizing the basal area of contact with the substratum (Gili et al. 1995). Clingers, such as Toucasia, colonized calm to moderate energy environments with only episodic sedimentation. In this muddy environment, most of the facies are characterized by wackestone textures. A common association of rudists with miliolids (and also diversified mono- and biserial foraminifera) is observed only in purely calcareous intervals and implies shallow-water depositional environments, possibly with raised sea-water salinity. The sediments suggest low-nutrient input, and evaporation probably slightly exceeded recharge

by rainfall, rivers or open ocean waters. This reflects a period of low-energy conditions and the few occurrences of hummocky cross-stratifications in the most proximal deposits emphasize quiet-water settings with occasional higher energy conditions. Conversely, the grainy, rudist- and milioliddominated facies indicate lower trophic conditions (oligotrophic). Changes in trophic level and clay input might be related to variable humidity patterns during sea-level rise and fall (Pittet et al. 2002). From a palaeobiological point of view, early Aptian rudist faunas show a European character whereas late Aptian faunas tend to be more cosmopolitan in Mediterranean regions (Masse & Gallo Maresca 1997; Masse et al. 1998). Apart from the main mid-Aptian turnover (Masse 1995) significant assemblages are recognized in the latest lower Aptian (Bedoulian) and within the Gargasian (early late Aptian). The mid-Aptian evolutionary changes in rudist faunas correspond to the onset of Radiolitidae and the extinction of the Caprinidae (Fenerci-Masse et al. 2006). Masse & FenerciMasse (2008) propose thermal shifts between cool and warm episodes fluctuating from 3 –5 8C, and the thermal threshold to explain the control of the demise of western European taxa in the range of 18 –20 8C. These temperatures are those controlling the geographic distribution of modern coral reefs, or the boundary between tropical and subtropical seas. Masse & Fenerci-Masse (2008) argue that rather than considering rudists in general as tropical, some of them may have been both tropical and subtropical and others tropical. They show that the regional disappearance of Lovetchenia, Agriopleura and Pachytraga is not linked to platform or rudist community demise but to cooling events documented by oxygen isotope studies. Some rudists seem to be more sensitive than others to the thermal regime (Masse & Fenerci-Masse 2008). The low diversity of rudists (mainly Requienid) in the lower part of the upper Aptian fits with the expected variation of the specific diversity of the Lower Cretaceous rudists of SE Spain following the mid-Aptian global event (Masse et al. 1998). In summary, the Aptian flora-faunal associations in the NeoTethys were probably triggered by both sea-level fluctuations and climatic changes affecting notably nutrient flux and seawater temperature. Both, early transgressions and humid phases caused an increased supply or mobilization of detrital components, and created mesotrophic to eutrophic conditions in which opportunistic species like orbitolinids thrived. This hypothesis is supported by the increasing wood-fragment content in siliciclasticrich Benasal Formation, but could also (partially) explain the absence of rudists within the deposits of Sequence I-1 whereas they flourished in Oman, Vercors, Provence and Helvetic platform domains

CRETACEOUS PLATFORM MARGIN IN SPAIN

during the early Aptian period (Fig. 13). During highstand, however, siliciclastic supply was cut off from the platform top, and/or possibly reduced through a lower run-off, and an oligotrophic environment existed, which was advantageous for the development of corals and rudists. Microbial facies is a typical dominant component in the early part of the early Aptian on the platform top, and it continued to dominate along the platform margins towards the end of the early Aptian (Spain and Oman plate margin).

Conclusions (1)

(2)

(3)

(4)

(5)

Four orders of sequences are defined (second to fourth) in excellent and continuous exposures of an Aptian carbonate-siliciclastic platform to basin transition in the Galve Basin in Spain (380 m thick, along a 16 km transect). The Aptian marine succession is subdivided into two large scale sequences separated by a middle Aptian sub-aerial exposure surface. At the base and at the top of the succession facies are siliciclastic continental to coastal plain. At the third-order scale a characteristic evolution of the flora-faunal fossil assemblages is observed, starting with orbitolinid-ooid dominated ramps in Sequence I-1, to a coral-stromatoporoidmicrobialite-dominated platform in Sequence I-2, to a rudist-dominated platform top in Sequence II-1, and finally a second episode of orbitolinid-ooid-dominated ramp system in Sequence II-2. A high quality carbon-isotope curve was constructed covering the full Aptian succession. This curve fits well other published curves measured in both benthic and pelagic domains. It provides a rare example of a complete Aptian curve in a platform to basin transition. The facies evolution in the Galve Basin resembles more the Arabian Plate pattern than the thin, short-lived earliest Aptian platforms along the northern NeoTethys margins (Provence, Vercors, Switzerland). Changes from orbitolinid-dominated to coral-stromatoporoid-microbialite and rudistdominated facies is interpreted as the result of changing depositional environments, from mesotrophic to oligotrophic, controlled by sea-level fluctuations and possibly associated with climatic variations affecting the amount of run-off.

The authors would like to thank J.-P. Masse and J. Garland for sharing their knowledge of the Urgonian carbonate platforms and stimulating discussions. V. Rommevaux is

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also acknowledged for his precious contribution to the preparation of the thin sections. Comments from two anonymous reviewers helped to improve the manuscript.

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CRETACEOUS PLATFORM MARGIN IN SPAIN Formations: Implications for the Early Cretaceous evolution of the Arabian Gulf Region. AAPG Bulletin, 80, 647–662. VAN B UCHEM , F. S. P., P ITTET , B. ET AL . 2002. Highresolution sequence stratigraphic architecture of Barremian/Aptian Carbonates systems in northern Oman and the United Arab Emirates (Kharaib and Shu’aiba Formations). GeoArabia, 7, 461–500. VAN B UCHEM , F. S. P., A L -H USSEINI , M., M AURER , F., D ROSTE , H. & Y OSE , L. 2010. Aptian Stratigraphy and Petroleum Habitat of the Eastern Arabian Plate. GeoArabia Special Publications, 4. VAN W AGONER , J. C., P OSAMENTIER , H. W., M ITCHUM , R. M., V AIL , P. R., S ARG , J. F., L OUTIT , T. S. & H ARDENBOL , J. 1988. An overview of the fundamentals of sequence stratigraphy and key definitions. In: W ILGUS , B. S. H. C. K., K ENDALL , C. G. ST . C., P OSAMENTIER , H. W., R OSS , C. A. & VAN W AGONER , J. C. (eds) Sea Level Changes: An Integrated Approach. SEPM Special Publications, 42, 39– 45. V ENNIN , E. & A URELL , M. 2001. Stratigraphie se´quentielle de l’Aptien du sous-bassin de Galve´ (Prov. De Te´ruel, NE de l’Espagne). Bulletin de la Socie´te´ Ge´ologique de France, 172, 397– 410. V ENNIN , E., S ORIA , A. R., P RE´ AT , A. & M ELE´ NDEZ , A. 1993. Ana´lisis secuencial durante el intervalo Barremiense-Aptiense en la Cubeta de Oliete. Cuadernos de Geologı´a Ibe´rica, 17, 257– 283. V ILAS , L., M ASSE , J. P. & A RIAS , C. 1995. Orbitolina episodes in carbonate platform evolution: the early Aptian model from SE Spain. Palaeogeography, Palaeoclimatology, Palaeoecology, 119, 35–45. W ALTER , B. 1991. Changement des faunes de bryozoaires dans le Valanginien supe´rieur des Alpes-de-HauteProvence. Paralle´lisme avec la crise observe´e dans le Jura a` la meˆme e´poque. Cretaceous Research, 12, 597–606. W EISSERT , H. 1989. C-isotope stratigraphy, a monitor of paleoenvironmental change; a case study from the Early Cretaceous. Surveys in Geophysics, 10, 1 –61. W EISSERT , H. & B REHERET , J. G. 1991. A carbonateisotope record from Aptian-Albian sediments of the Vocontian Trough (SE France). Bulletin de la Socie´te´ Ge´ologique de France, Huitie`me Serie, 162, 1133–1140. W EISSERT , H. & L INI , A. 1991. Ice age interludes during the time of Cretaceous greenhouse climate.

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High-resolution seismic stratigraphy of the Shu’aiba and Natih formations in the Sultanate of Oman: implications for Cretaceous epeiric carbonate platform systems H. DROSTE Shell Technology Oman, P.O. Box 74, PC 116 Mina Al Fahal, the Sultanate of Oman Present address: Shell International Exploration and Production B.V., P.O. Box 60, 2280 AB Rijswijk, The Netherlands (e-mail: [email protected]) Abstract: The Shu’aiba and Natih Formation carbonates are important hydrocarbon reservoirs in the Sultanate of Oman. They consist of stacked deepening and shallowing-upward depositional cycles within an extensive middle Cretaceous epeiric carbonate platform. Geological models for these units traditionally assume a layer-cake stratigraphy and a high lateral continuity of facies. This is based on the assumption that epeiric platforms consist of ramps with low depositional gradients, broad facies belts and gradual facies transitions. However, high-resolution 3D seismic data covering large areas of the platform have revealed a more complicated internal stratigraphic architecture and have led to a new geometrical model for these epeiric platform systems. The transgressive part of the cycles is dominated by a low angle ramp depositional profile with localized development of algal-dominated mounds. Differential carbonate growth led to a topography of shallow carbonate shoals and intra-platform ‘basins’ with water depths reaching several tens to 100 m. During the regressive part of the cycles these basins were progressively filled-in by prograding carbonate rudist shoal complexes with depositional slopes of 0.58 to more than 308. Clinoform belts on seismic show a wide range of progradational geometries ranging from closely spaced, laterally continuous ‘tramlines’ to irregular wedges and noses. The cycle tops are characterized by bedrock incisions and the influx of fine clastic sediments that fill in remnants of the intra-platform basins. The seismic images show that previous stratigraphic models for these carbonates oversimplified and flattened the stratigraphy. As a result the stratigraphic trapping potential and the internal reservoir heterogeneity for these systems has been underestimated.

Epeiric carbonate platforms are a category of carbonate platforms that were deposited in epicontinental shallow seas and that are characterized by their enormous laterally extent (100 –1000s km wide, see Fig. 1). These platforms are very common in the geological record and host some of the world’s richest petroleum systems, such as the Permian to Cretaceous of the Middle East (Murris 1980). Unfortunately there are no modern analogues for these platforms and current stratigraphic models of these carbonate systems are mainly based on the interpretation of subsurface well data. Epeiric platforms are considered to have negligible topography and very low angle depositional slopes of less than 18 (Irwin 1965). Deposition on these platforms is dominated by cycles of low energy shallow subtidal (water depths less than 10 m) to intertidal carbonates occurring in very wide facies belts (Irwin 1965; Tucker & Wright 1990). In places on the platform, muddy sedimentation occurs within deeper water basins, commonly under anoxic bottom-water conditions (Murris 1980). Depositional cycles on the platform are

widely correlatable and have sheet-like geometries. This is thought to be related to the flat depositional profile and high sedimentation rates which allows rapid progradation after each slight relative sea-level rise. Though these sheets represent diachronous deposition, the biostratigraphic resolution is not good enough to prove this with stratigraphic dating. Depositional profiles on the epeiric platforms are considered to be of very low angle and a carbonate ramp depositional model is commonly used to describe platform sedimentation (e.g. Burchette & Britton 1985; Burchette 1993; van Buchem et al. 1996, 2002a; Davies et al. 2002). The ramp model assumes a gently sloping surface on which shallowwater facies pass gradually into deeper water and then into basinal sediments (Tucker & Wright 1990). However, seismic data suggest that the internal architecture is more complex than a simple ramp system (Droste & van Steenwinkel 2004). The seismic data show a variety of depositional architectures: differentiated platforms with clear breaks in the slopes, mounded features and clinoforms with steeper angles than that considered

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 145–162. DOI: 10.1144/SP329.7 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Main carbonate platform types (modified from Tucker & Wright 1990).

typical of ramps. These features have important implications for our understanding of the distribution and internal architecture of hydrocarbon traps and reservoirs within these systems and hence the need to revisit previous interpretations. In order to understand the stratal geometries of the epeiric platform systems better, a regional seismic review of the Cretaceous was carried out in north Oman in a study area covered by extensive merged 3D seismic projects made available by Petroleum Development Oman (Fig. 2). The seismic

Fig. 2. Location map showing study area.

data were scanned for stratal geometries using both flattened vertical sections and seismic attribute maps. The seismic observations were integrated with the lithological and stratigraphic information obtained from wells. This resulted in a better understanding of the controlling factors of the internal architecture and evolution of the depositional cycles on this epeiric carbonate platform. The objective of this paper is to present a geometrical refinement of the model for epeiric platform carbonate deposition based on the integration of seismic and well data. The low depositional profiles invoked for some epeiric systems is likely to be an artefact resulting from long distance correlation of 1D well data or outcrops which tend to flatten the stratigraphy. Two examples from the Cretaceous epeiric carbonate platform in Oman, the Shu’aiba Formation and the Natih-e Member are used to illustrate the internal architecture. Each example shows how the seismic observations changed the previous sequence stratigraphic models that were based on well and outcrop data.

Geological setting The Shu’aiba and the Natih Formation in Oman are part of a large Cretaceous carbonate platform system (Fig. 3; e.g. Hughes-Clarke 1988; Pratt & Smewing 1993; Droste & van Steenwinkel 2004). This platform started off as a strongly prograding Early Cretaceous carbonate shelf that prograded some 250 km over a tilted, truncated and subsequently drowned Middle Jurassic carbonate platform. Shelf progradation ceased in Hauterivian/ Barremian times after which the margin shows an overall aggradational growth pattern, probably in response to a strongly rising relative sea-level. Flooding of the relatively flat area landward of the margin resulted in the formation of a shallow epeiric shelf that extended some 1000 km from the margin to the exposed Arabian Shield in the SW. This epeiric shelf persisted for some 40 Ma until the Turonian when it was terminated by a regional uplift preceding the obduction of the Oman ophiolite along the northern margin (Glennie et al. 1974; Loosveld et al. 1996). Carbonate deposition predominated in this epeiric sea and in total some 700 m of mainly shallow-water carbonates accumulated. These are internally organized into regional correlatable deepening and shallowing-upward packages that were driven by higher order sea-level changes superimposed on the overall increase in accommodation. Thickness of these cycles ranges from several tens to 150 m and maximum water depths are estimated to be in the order of some 80 m (van Buchem et al. 2002a, b; Droste & van Steenwinkel 2004).

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Fig. 3. Geological cross section through the North Oman Cretaceous Carbonate platform, see Figure 2 for location of section (Modified from Droste & van Steenwinkel 2004).

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Carbonate deposition on the platform was repeatedly interrupted by sub-aerial exposure and influx of fine-grained clastic sediments, mainly clays, sourced from the exposed Arabian shield in the south and fringing exposures of Palaeozoic sediments in the SW. The leached carbonates and overlying clays form a number of reservoir-seal pairs that now host important hydrocarbon accumulations. Both the Shu’aiba Formation and the Natih-e Member discussed in this article each represent a deepening and shallowing-upward event on the platform and are major hydrocarbon reservoirs in the Sultanate of Oman.

Example 1: Shu’aiba Formation The Shu’aiba Formation is the uppermost unit of the Kahmah Group (Hughes Clarke 1988) and consists of an Aptian carbonate complex of up to 130 m thick (Fig. 4). It is separated from the clean carbonates of the underlying Kharaib Formation by an exposure surface, which is overlain by an easily recognizable marker bed of tight limestones (Hawar Member). It is unconformably overlain by a thick shale interval of the Nahr Umr Formation which forms an excellent seal and explains the prolific occurrence of hydrocarbons in the Shu’aiba.

Fig. 4. Type logs for the Shu’aiba Formation in NW and Central Oman, see Figure 5 for well location.

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Fig. 5. Palaeogeography for the Shu’aiba Formation at the end of the Early Aptian (plotted on present day orientation).

In northwestern Oman (area around well A on Fig. 5), the Shu’aiba can be subdivided into a Lower and Upper Member (Hughes Clarke 1988; Mohamed et al. 1997, see also Fig. 4). The Upper Member is Late Aptian in age (Witt & Go¨kdag

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1994; van Buchem et al. 2002a) and consists of interbedded calcareous clays and limestones. The Lower Member is Early to Late Aptian in age (van Buchem et al. 2002a) and consists of clean limestones. The Upper Member corresponds with what has been defined as the Bab Member in the United Arab Emirates (UAE, Hassan et al. 1975; Hughes Clarke 1988). This represents the infill of an intrashelf basin (‘Bab basin’), which covered large parts of the UAE and extended into northwestern Oman (Fig. 5, Murris 1980; van Buchem et al. 2002a). Outside the area of the Bab basin, the Shu’aiba consists only of lower Aptian limestones and no subdivision into members is applied. A regional sequence stratigraphic model for the Shu’aiba has been proposed by van Steenwinkel (1992) and Witt & Go¨kdag (1994) based on subsurface transect in Central and North Oman, and by van Buchem et al. (2002a), based on outcrop and subsurface transects in North Oman and the United Arab Emirates. These models present a well-based sequence stratigraphic interpretation of the Shu’aiba using well logs calibrated with all available core data and biostratigraphic information. The model by Van Steenwinkel (1992, Shell internal publication) is illustrated in Figure 6. The correlation uses marine flooding surfaces as isochronous correlation markers and patterns in the stacking of the sedimentary cycles. A major flooding surface has been recognized based on a gamma-ray peak/trend change and maximum landward extent of deepest-water facies consisting of pelagic mudstones. Below the maximum flooding surface, the lower part of the

Fig. 6. Well-based sequence stratigraphic framework for the Shu’aiba Formation illustrated in a section from NW to SE (Van Steenwinkel 1992, Shell internal publication). See Figure 6 for location map.

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Shu’aiba is shown as back-stepping, low-angle landward onlapping wedges dominated by algal (Bacinella-Lithocodium) facies lacking rudists overlying lagoonal mudstones of the Hawar Member. Above the maximum flooding surface, the stacking patterns are aggradational to overall progradational (up and out-building) to toplapping and consists mostly of rudist shoal facies. It is also suggested that the base of the argillaceous ‘Upper Shu’aiba’ (Bab Member) in the Lekhwair area represents the correlative conformity that corresponds to the top Shu’aiba unconformity in the south. It is implied that the Upper Shu’aiba/Bab Member is the lowstand (or restricted highstand) that followed the Shu’aiba highstand in the south. The correlation strategy used for the interpretation follows all the correct geological principles (see Homewood et al. 2000). Walther’s law of correlation of facies has been applied, time-lines follow a depositional profile from shallow to deeper water facies and the model is consistent with all the data used. The cross-section shows clear stratal geometries such as onlap, downlap and truncation. However, the proposed inclined geometries are extremely flat: the boundary of the Bab basin has a dip of only 0.38; the suggested dips of the downlapping surfaces are less than 0.018. A regional review of 3D seismic surveys from different parts of Oman in the present study showed a much more complex stratigraphic architecture (Fig. 7). In several areas, mounded geometries occur within the Shu’aiba. The base of the mounds is located just above the top Kharaib/base Shu’aiba reflector and the top coincides with the top Shu’aiba. The slopes of the mounds are very low, about 18, and are onlapped by other intra-Shu’aiba reflectors. Seismic mapping of the mounds shows that they have a very irregular outline forming a complex of patches and ridges up to a few kilometres wide. Core data show that these mounds consist of algal Bacinella/Lithocodium facies and Orbitolinids, although rudists are lacking. In many areas, clinoforms have been observed, in places clearly prograding away from mound complexes. The clinoforms occur in the upper part of the Shu’aiba and the dip calculated from the seismic varies from 0.58 to 58, which is much steeper than suggested in correlations based on well data alone. Dipmeter data from wells suggest that locally dips of up to 358 are reached. Core data show that the composition of these clinoforms can range from very coarse-grained rudist debris to fine-grained carbonate sands or marls in the Bab basin. The higher clinoform angles are associated with coarser sediment debris and predominantly occur in the Lower Shu’aiba. These clinoforms show strong lateral variation in the amount and direction of

progradation. This points to localized input of sediment, probably from nearby rudist biostrome patches. The more marly clinoforms of the Upper Shu’aiba form laterally continuous parallel belts that can be followed over more than 100 km (Pierson et al. 2010). These may suggest that alongshore transport played an important role in the distribution of these sediments. The seismic data added significantly more detail to the understanding of the depositional patterns during the overall transgressive-regressive cycle than had been recognized from well data alone. Figure 8 shows a cross-section through the Shu’aiba platform with correlations guided by both seismic and well-derived data. During the transgressive phase, a strong differentiation in bathymetry on the platform is suggested from the development of algal mounds. Between these mounds fine-grained, nannoconid-dominated, in places organic-rich, deeper-water sediments were deposited. These mounds were colonized by rudist shoal complexes and through progradation developed into larger platforms. The clinoforms of this progradational phase are much steeper and more multidirectional than suggested by the well log correlations. The deeper water areas between the platforms were filled progressively by the clinoform belts, which record increasingly offlapping geometries and indicate relative falling sea-level. As sea-level continued to fall, increasing amounts of fine-grained siliciclastic sediments, shed from the Arabian Shield, entered the intra platform basins and argillaceous carbonates and carbonate-rich claystones infilled the low angle clinoform complexes. The overall sequence stratigraphic patterns based on regional well data shown in Van Steenwinkel (1992, Shell internal publication) and van Buchem et al. (2002a) are confirmed by the seismic data; however the seismic shows that the internal geometries are much more complicated. Within the prograding complexes, even when the clinoform angles are relatively low, lateral continuity of individual sedimentary units is limited, generally in the order of a few hundred metres to a few kilometres in dip direction, and little can be correlated between wells, even on a field scale.

Example 2: Natih-e member The Natih-e is one of the seven members of the Natih Formation, which consists of mud-supported and some grain-supported limestones with local rudist development, alternating with calcareous clays (e.g. Hughes Clarke 1988; van Buchem et al. 1996). The Natih-e is Late Albian to Early Cenomanian in age and reaches a thickness of about 150 m (see Fig. 9). It consists of a thin clay interval at the base overlain by a thick carbonate package.

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Fig. 7. Examples of seismic stratal geometries in the Shu’aiba Formation. (a) Mounded features that started to form at or very close to the base Shu’aiba. (b) Amplitude map showing outline of the mound complexes represented by high amplitudes (courtesy O. Al-Jaaidi, PDO). (c) Instantaneous phase line showing clinoform complex in central Oman. See Figure 5 for location map.

152 H. DROSTE Fig. 8. Examples of stratigraphic correlations of the Shu’aiba Fm that honour the observed seismic geometries: (a) along the Bab basin margin, and (b) in the interior platform. With well data alone, it is not possible to capture the inclined geometries in the upper part of the Shu’aiba. See Figure 5 for location map.

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Fig. 9. Type log for the Natih-e member, for location see Figure 10 (modified from Droste & van Steenwinkel 2004).

Organic-rich chalk levels with source rock potential occur in the basal part of the carbonate unit (van Buchem et al. 2002b; Droste & van Steenwinkel 2004). The top of the e member is defined at the base of the thick clay interval of the basal Natih-d member. The whole member represents one deepening and shallowing-upward cycle on the epeiric shelf associated with the creation and subsequent infill of an intra-shelf basin (van Buchem et al. 1996, 2002b; Droste & van Steenwinkel 2004, see Fig. 10). The Natih-e depositional cycle is the equivalent of the Mauddud, Shilaif and Mishrif succession in the southern Arabian Gulf (Burchette 1993). van Buchem et al. (1996) presented a stratigraphic model for the Natih-e Member in North Oman based on high-resolution stratigraphic correlation of outcrops in the Oman Foothills (see Fig. 11). Schwab et al. (2005) constructed a synthetic seismic section through the Natih Formation based on this transect. The correlation in Figure 11 shows that the Natih-e Member comprises one major transgressiveregressive cycle. During the transgression, an aggradational trend developed, with the construction of a shallow-water carbonate platform and the deposition of organic-rich limestone in the adjacent intrashelf basin. The regressive part shows a clear progradation of shallow-water carbonate platform deposits. Higher-order cycles are superimposed on the overall trend. Like the Shu’aiba example discussed above, the correlation approach that has been used in this study is geologically correct, however, the resulting depositional profile is

Fig. 10. Palaeogeography for the latest Albian Lower Natih-e Member. Outline in the Gulf area (modified from Murris 1980).

extremely flat: the maximum dips suggested on the diagram are less than 0.18 (Fig. 11). Seismic data however reveal that the suggested depositional profile is too flat and that internal geometries are far more complex (Fig. 12). Clear inclined geometries can be recognized that show different directions and delineate prograding carbonate platform systems, separated by (clay-filled) intra-platform basins. Dip angle of the clinoforms (between 0.58 and 58) is related to the grain fabric. Low-angle clinoforms occur in the marly units and high-angle clinoforms are composed of platformderived grainstones. Finer-grained clinoform complexes form straight, parallel, laterally continuous belts (.100 km), which may suggest that alongshore transport played an important role in the distribution of these sediments. The coarser grained clinoform complexes show very irregular progradation patterns and are laterally less continuous (few tens of km). This may reflect localized input of carbonate sediments and limited alongshore transport of the coarser grains. Mapping of the clinoform trends shows that the Natih-e consists of a number of carbonate platforms that prograde into, sometimes completely filling, intervening basins and seaways (Droste & van Steenwinkel 2004). The dips of the clinoforms indicated by seismic data are much steeper than those suggested by the outcrop correlation in Figure 11.

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Fig. 11. Outcrop correlation model for the Natih-e in the Oman Foothills area, see Figure 10 for location map (modified after van Buchem et al. 1996).

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Fig. 12. Seismic amplitude map of top Natih-e reflector in Central Oman showing several carbonate platforms with clinoform belts. The displayed seismic line shows the clinoform geometries from the platform interior to the margin of the intra platform basin platform. For location map see Figure 10.

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To date, well-defined mounded-like geometries on seismic, as those described above for the Shu’aiba have not been recognized in the Natih-e. However, some mounding and aggradational stacking has been recognized in a field in Central Oman, which is located above a salt dome (Masaferro et al. 2003, 2004). Cores show that the Natih-e in this area is composed of algal and Orbitolina dominated sediments (Smith et al. 2003). In the upper part of the Natih-e a network of channels has been observed on seismic attribute maps, which may be more than 1 km wide and reach lengths of over 30 km (Fig. 13). Similar channels have been observed in outcrops of the Natih-e in North Oman (van Buchem et al. 1996; Gre´laud 2005; Gre´laud et al. 2006). These channels are bedrock incisions of up to 25 m deep and up to 2 km wide that were formed during sub-aerial exposure of the platform following a relative fall in sea-level. The incisions were filled during the subsequent relative rise in sea-level. The fills are made up of several genetic units, with a variety of carbonate facies and smaller scale channel forms that are dolomitized in places. The channels drained into the intra shelf basins and seaways between the platforms. Infill of these lows on the shelf during the sea-level low stand

was a mixture of erosional products from the exposed platforms and clastic sediments from the Arabian Shield. In Oman, the clastic sediments are predominantly clays (Droste & van Steenwinkel 2004), in other areas more proximal to the clastic source coarser-grained clastic sediments filled these basins as illustrated by Azzam (1994). Figure 14 shows a well correlation panel along the seismic section shown in Figure 12, superimposed with the observed clinoform dips. The clays at the base and top of the Natih-e cycle are regionally correlatable marker beds indicating a flat depositional topography. However, seismic data shows that this is not the case for the carbonates in between. Though the overall dips observed on seismic are very small: the ‘high-angle’ dips are only 28, whereas the low-angle dips are less than 0.58, these are much steeper than those suggested by well correlations alone. Despite the gentlest dips, wells that are a few kilometres apart (in dip direction) are hardly correlatable.

Discussion The examples illustrate how stratigraphic models of these platform interior carbonate systems, based on the correlation of 1D well and core data

Fig. 13. Channel incision into top of Natih-e imaged on an azimuth map of the top Natih-e seismic reflector. For location see Figure 10.

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Fig. 14. Well correlation panel for the Natih-e along the seismic line shown in Figure 12 with a projection of the seismically defined clinoforms. Compared to the outcrops correlation shown in Figure 11, the seismically constrained interpretation shows much stronger downlapping geometries.

oversimplify and flatten the stratigraphy, even between closely-spaced wells. This is especially true in the upper, shallowing-upward part of the cycles. The reason is that we always try to push markers/log picks through as far as possible and also that correlations are made with a very low angle ramp profile in mind. The correlation in these systems is further complicated by the fact that the dips involved are still relatively low and cannot be recognized as such in cores or on image logs. Seismic data is essential for recognizing these stratal geometries. In many cases, the quality and/or the resolution of the seismic may not be good enough to see all the stratal geometries in detail. For example, with increasing depth the frequency on the seismic signal will decrease and progressively fewer stratal features can be resolved which may even lead to misinterpretations (see Vail et al. 1977; Stafleu et al. 1994). In cases where vertical sections show no clear or very subtle stratigraphic variation, seismic attributes displayed in map view (using 3D seismic) still can reveal detailed depositional features and trends. Input from outcrop studies, calibration with well log and core data and understanding of the regional stratigraphic framework is essential to recognize and interpret the seismic features. The Shu’aiba and Natih-e examples show that there was a common evolution of depositional patterns and stratigraphic architecture on the Cretaceous epeiric shelf in response to major relative sea-level rises and falls which can serve as a reference for depositional geometries that might be expected in other epeiric carbonate systems. This model is described below and illustrated in

Figure 15. This seismically-derived model confirms the overall sequence stratigraphic framework presented in earlier well log and outcrop-based studies, but provides significantly more geometrical detail. A regional sequence boundary is present at the base of the sequence exhibiting sub-aerial exposure of the underlying platform carbonates. In the case of the Shu’aiba, this boundary is a major regional erosional surface associated with sub-aerial exposure separating open shallow-marine carbonates of the Upper Kharaib from intertidal deposits of the Hawar Member (van Buchem et al. 2002a; Droste 2007). The presence of higher-order cycles makes it more difficult to assign the base of Natih-e cycle. van Buchem et al. (2002b) picked this boundary at a hardground surface developed at the top of the Natih-g member as this marks the onset of the development of a differentiated topography. Regional well log correlations show that during Natih-g times, a differentiated topography was present (Droste & van Steenwinkel 2004). An alternative (also suggested in van Buchem et al. 1996) and preferred boundary proposed here is the top of the Natih-f Member, which shows possible soil development and a well-developed bored hardground surface (Fahud Field, Davies et al. 2002; Internal PDO report). Along the oceanward margin of the platform, lowstand shelf margin wedges were deposited while the epeiric platform was exposed. The base Shu’aiba lowstand is exposed in outcrops along the northern edge of the Oman Mountains. It consists of a wedge of gravel beaches and microbial build-ups grading downslope into cross-bedded rudistid bioclastic grainstones and, locally, microbialites (Hillga¨rtner et al. 2003). The platform

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Fig. 15. Geological model for a depositional cycle on the Cretaceous epeiric carbonate platform in Oman in response to, from top to bottom, a rise and fall in relative sea-level.

margin equivalent of the Natih Formation was situated oceanward of the Shu’aiba wedge. This interval is now located below the ophiolite nappes on the northern side of the Oman Mountains and is not exposed at the surface nor penetrated by wells. During the initial rise in sea-level the platform was progressively flooded and the clastic sediments were pushed back towards the Arabian Shield (Fig. 15a). Low depositional profiles resulted in laterally extensive units of marginal to shallow-marine sediments, typically dominated by algal facies. A relatively high influx of fresh water with a high nutrient input may have favoured the algal-dominated carbonate factory. Over large parts of the shelf, the carbonate factory was not able to keep up with the rise in sea-level resulting in a regional differentiation with local drowning of the epeiric shelf and a deepening upward facies trend, which can be observed both in the basal, Shu’aiba and Natih-e. Locally, carbonate sedimentation managed to keep up with rising sea-level, which resulted in the development of algal mound complexes (light blue areas on Fig. 15b). Location of these mound complexes may have been controlled by subtle elevation differences (few metres maximum), for example, erosional high or structural elevation related to faulting or salt doming or flexural isostacy (Aigner et al. 1989). With rising sea-level a strongly differentiated topography developed with water depths ranging from 0 to some 100 m. In the Shu’aiba

mound complexes developed during the Early Aptian; these have been recognized both on seismic and in well data from Oman (see above) and the United Arab Emirates (Yose et al. 2006; Hughes 2000). The Natih-e mounds developed during the Late Albian but as this interval has a low hydrocarbon potential over most of the country very little petrographical data is available on this facies apart from one field in Central Oman (Masaferro et al. 2003, 2004; Smith et al. 2003). Analogues for the type of mounds that can be expected in the Natih-e could be the Upper Albian algal mounds that have been reported from northern Spain (Garcia-Mondejar & FernandezMendiola 1995). Prolific nannoplankton occurred in the open marine waters on the shelf and are recorded in the pelagic carbonates (chalks) between the mound complexes. In the deeper parts of the shelf, these chalks are laminated and contain high amounts of organic matter, which may indicate anoxic bottomwater conditions. The total organic matter content (TOC) for the Shu’aiba in Oman may reach up to 4% TOC (van Buchem et al. 2002b) whereas the richest Shu’aiba source rocks occur in the central part of the Bab basin offshore the United Arab Emirates with TOC values up to some 10% and a total thickness of some 30 m (Taher 1997). The TOC of the pelagic sediments of the Natih-e in Oman rarely exceeds 5% (Terken 1999). Up to 14% weight TOC have been reported from the more

SEISMIC STRATIGRAPHY OF CRETACEOUS PLATFORMS

basinal equivalents of this interval in Saudi Arabia (e.g. the Safaniya Member of the Wasia Formation; Newell & Hennington 1983). The age of the organic-rich intervals suggests that they were linked to worldwide anoxic events (Schlanger & Jenkyns 1976). The Early Aptian Shu’aiba interval is time equivalent to the oceanic anoxic event 1a or Selli event (OAE-1a), which is recognized in many carbonate platforms worldwide (Erbacher & Thurow 1997; Weissert et al. 1998). The Natih-e organic-rich interval may correspond to the widely distributed Late Albian Oceanic Anoxic Event 1d (OAE-1d) or Breistroffer Event (Bornemann et al. 2005; Wilson & Norris 2001). Preservation of large amounts of organic carbon during these events has been linked to productivity changes caused by variations in nutrient supply and/or the presence of a stratified water column with low bottom-water oxygenation. During the sea-level highstand, the carbonate factory in both in the Shu’aiba and the Natih-e was dominated by a grainy facies, rich in rudists (indicated by dark blue on Fig. 15c). In the case of the Shu’aiba, the change in carbonate factory has been related to changes in trophic and energy levels on the shelf (Pittet et al. 2002; van Buchem et al. 2002a). The algal-dominated carbonate factory was characterized by low energy and mesotrophic conditions, whereas the rudist factory was controlled by higher energy and oligotrophic conditions. The nutrient levels may have changed in response to climatic changes linked to the OAE-1a event (Pittet et al. 2002; Hillga¨rtner et al. 2003; Immenhauser et al. 2005; van Buchem et al. 2002a) or may simply reflect a relative decrease of fluvial influx following the major sea-level rise. Similar conditions, now linked to the OAE-1d event, may have played a role in the case of the Natih-e. Rudistid shoal complexes preferentially developed in the shallow waters above the algal mounds. Prolific growth of the rudists and reworking of the rudist biostromes by storms provided abundant sediment, which was shed into deeper water around the shoals. These redeposited shallow-water carbonates formed prograding slope complexes leading to a progressive growth of the mound complex into small platforms. With the decreasing rate of relative sea-level rise these grew and eventually coalesced into larger platforms covering most of the shelf. As relative sea-level began to fall these platforms became sub-aerially exposed (Fig. 15d). A clear erosional topography developed as shown by the channel incisions within the top of the Natih-e member (Gre´laud 2005; Gre´laud et al. 2006). An erosional topography is also indicated by metrescale thickness variations between closely spaced wells in the Shu’aiba. A further indication for significant erosion is the absence of shallow-water

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Shu’aiba facies below the Nahr Umr unconformity over large areas in Northern Oman where the Albian Nahr Umr shales overlie Lower Aptian outer to middle ramps sediments. Both the channel incisions and the facies break suggest an erosional topography in the order of several tens of metres. Though a tropical humid climate is suggested by the presence of coal layers in well penetrations of the Natih in North Oman and the United Arab Emirates and the occurrence of silicified wood fragments in the Shu’aiba in the Huqf area (Platel et al. 1992), evidence for significant meteoric diagenesis related to sub-aerial exposure is scarce (Vahrenkamp 1996; Immenhauser et al. 2000). Wave erosion during the subsequent flooding of the exposed platform may have removed most of the weathered top of the platforms. Deep cutting leaching that is clearly related to exposure is limited to the coarser grained, more permeable, rudist-bearing sediments as is suggested by infiltration of the Nahr Umr shales up to some 40 m into the underlying Shu’aiba Formation. Increased fluvial input from the Arabian Shield during the lowstand led to the deposition of clastic sediments in the remaining seaways between the exposed platforms. Carbonate production in these seaways is progressively smothered by clastic sediments leading to a mixed carbonate-clastic, and eventually, a clastic infill. At the end of the cycle, the whole epeiric shelf may have been sub-aerially exposed. During this time the platform was incised by river systems carrying clastic sediments from the Arabian Shield to the shelf margin and into the open ocean. On the external parts of the epeiric shelf the end product of one relative sea-level cycle is a carbonate layer, internally composed of a deepening and shallowing-upward cycle, bounded by exposure surfaces and thin clastic interval. This cycle has a more or less flat parallel top and bottom that can be correlated over large distances but internally has a highly complex stratal architecture. A ramp type depositional profile like that present along the southern coast of the Arabian Gulf (see also Purser 1973), only applies to the lower transgressive part of the cycle. Most of the time the epeiric platform had a differentiated topography, with platforms, initially small and separated by intra-shelf basins, later covering most of the shelf and separated by seaways. The outline of these platforms and seaways resembles that of the present-day Bahamas. Internally, the Bahamas platforms also show a similar architecture with strong progradational geometries and small platforms merging into larger ones; however the vertical scale is much larger with water depths of many hundreds of metres (Eberli & Ginsburg 1987, 1989).

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The geometrical refinement of the sequence stratigraphic model discussed in the present paper may also apply to epeiric platforms of other ages both in the Middle East and other parts of the world. A key component of this model, the presence of intra-platform basins, is a common feature of large platform systems and has been reported not only for several other stratigraphic intervals on the Arabian Plate such as the late Oxfordian–early Kimmeridgian, early –mid Valanginian and the mid-late Barremian (e.g. Murris 1980; Droste 1990) but also in Cambrian, Devonian and Mesozoic platforms in North America (Eliuk 1978; Markello & Read 1982; Read 1985; Elrick 1996). Some of these basins may have formed as result of structural related variations in subsidence. These basins are expected to show an increased sediment thickness compared to the surrounding platform. A fundamental stratigraphic observation is that intraplatform basins with a sediment infill that is more or less similar to the thickness of the adjacent platform carbonates indicates that differential carbonate growth has been the driving mechanism of the differentiation in topography.

Conclusions Seismic data of the Shu’aiba and Natih-e depositional cycles from the interior of Oman have added significantly more detail to the understanding of the internal stratigraphic architecture of this Cretaceous epeiric carbonate platform system. While confirming the sequence stratigraphic framework as defined by earlier studies based on outcrop and well data, the observations also show that the lateral heterogeneity on a scale of a few hundred metres to kilometres is much larger and more complex than has previously been assumed. Within the stacked deepening and shallowingupward packages that make up this platform, a common pattern can be recognized of differential vertical carbonate growth during the transgressive phase followed by lateral growth through progradation during the highstand and early lowstand phase. Development of algal microbial mounds during the transgression played an important role in building the differentiated topography as is clearly shown by seismic examples of the Shu’aiba Formation. Drowning of the platform interior during the transgression led to the development of intra-platform basins with restricted bottom-water circulation and deposition of organic-rich sediments with source rock potential. Infill of these basins was the result of lateral progradation of rudist-dominated carbonate shoals with slope angles varying between 0.58 and 358. The slope angles and lateral continuity of the clinoform belts were controlled by grain fabric

and sediment transport: laterally extensive low angle systems are composed of fine-grained sediment brought in by alongshore transport, high angle systems are irregular in outline, laterally less extensive and are associated with local input of coarse-grained sediments. The depositional cycle ended with a relative drop in sea-level that caused sub-aerial exposure associated with bedrock incisions of the platform top. During this phase the remnants of the intra-platform basins were filled with fine-grained clastic sediments derived from the hinterland. Comparison of well-based high-resolution stratigraphic correlations with seismic patterns suggest that in the former the low angle depositional slopes are not captured due to an over correlation of individual markers/log picks leading to an artificial flattening of the stratigraphy. Recognition of the complex internal stratigraphy described in this paper has important implications for hydrocarbon prospectivity and reservoir development in these epeiric platform carbonates. On a regional scale, it will help to define the stratigraphic trapping geometries. At the development scale, it will lead to a better understanding and prediction of reservoir heterogeneities that will affect flow through the reservoir and the ultimate recovery. When investigating the stratigraphic architecture of these epeiric carbonate systems, it is essential to scan available seismic data carefully as this may reveal kilometre-scale stratigraphic geometries that are not captured by well data. An assessment of the subseismic scale heterogeneities may be obtained from outcrop analogue studies. Stratigraphic models based on well data alone tend to underestimate the stratigraphic complexity of these carbonate systems. This study has been carried out in the Shell JVR Centre for Carbonate Studies at the Sultan Qaboos University of Oman. Petroleum Development Oman and the Ministry of Oil and Gas of the Sultanate of Oman are gratefully acknowledged for making available the subsurface data. This paper benefited from a thoughtful review by Frans van Buchem (Doha, Qatar) and an anonymous reviewer.

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S TAFLEU , J., E VERTS , A. J. W. & K ENTER , J. 1994. Seismic models of a prograding carbonate platform: Vercors, southeast France. Marine and Petroleum Geology 11, 514– 527. T AHER , A. A. 1997. Delineation of organic richness and thermal history of the Lower Cretaceous Thamama Group, East Abu Dhabi: a modeling approach for oil exploration. GeoArabia, 2, 65–88. T ERKEN , J. M. J. 1999. The Natih petroleum system of North Oman. GeoArabia, 4, 157–180. T UCKER , M. E. & W RIGHT , V. P. 1990. Carbonate Sedimentology. Blackwell Science, Oxford. V AHRENKAMP , V. C. 1996. Carbon isotope stratigraphy of the upper Kharaib and Shu’aiba formations: implications for the Early Cretaceous of the Arabian Gulf region. American Association of Petroleum Geologists Bulletin, 80, 647– 662. V AIL , P. R., T ODD , R. G. & S ANGREE , J. B. 1977. Seismic Stratigraphy and Global Changes of Sea-level: Part 5. Chronostratigraphic Significance of Seismic Reflections: Section 2. Application of Seismic Reflection Configuration to Stratigraphic Interpretation. In: P AYTON , C. E. (ed.) Seismic Stratigraphy Applications to Hydrocarbon Exploration. American Association of Petroleum Geologists Memoir, 26, 99–116. VAN B UCHEM , F. S. P., R AZIN , P. ET AL . 1996. High resolution sequence stratigraphy of the Natih formation (Cenomanian/Turonian) in northern Oman: distribution of source rocks and reservoir facies. GeoArabia, 1, 65–91. VAN B UCHEM , F. S. P., P ITTET , B. ET AL . 2002a. Highresolution sequence stratigraphic architecture of Barremian/Aptian carbonate systems in northern Oman and the United Arab Emirates (Kharaib and Shu’aiba Formations). GeoArabia, 7, 461– 500. VAN B UCHEM , F. S. P., R AZIN , P., H OMEWOOD , P. W., O TERDOOM , H. & P HILIP , J. 2002b. Stratigraphic organization of carbonate ramps and organic-rich intrashelf basins: Natih formation (middle Cretaceous) of northern Oman. American Association of Petroleum Geologists Bulletin, 86, 21–54. W EISSERT , H., L INI , A., F O¨ LLMI , K. B. & K UHN , O. 1998. Correlation of Early Cretaceous carbon isotope stratigraphy and platform drowning events: a possible link. Palaeogeography, Palaeoclimatology, Palaeoecology, 137, 189–203. W ILSON , P. A. & N ORRIS , R. D. 2001. Warm tropical ocean surface and global anoxia during the midCretaceous period. Nature, 412, 425– 429. W ITT , W. & G O¨ KDAG , H. 1994. Orbitolinid biostratigraphy of the Shu’aiba Formation (Aptian), Oman – implications for reservoir development. In: S IMMONS , M. D. (ed.) Micropalaeontology and Hydrocarbon Exploration in the Middle East. British Micropalaeontological Society Publication Series, Chapman & Hall, Cambridge, 221– 242. Y OSE , L. A., R UF , A. S. ET AL . 2006. Three-dimensional characterization of a heterogeneous carbonate reservoir, Lower Cretaceous, Abu Dhabi (United Arab Emirates). In: H ARRIS , P. M. & W EBER , L. J. (eds), Giant Hydrocarbon Reservoirs of the World: From Rocks to Reservoir Characterization and Modeling. American Association of Petroleum Geologists Memoir, 88, 173–212.

Channelized systems in an inner carbonate platform setting: differentiation between incisions and tidal channels (Natih Formation, Late Cretaceous, Oman) C. GRE´LAUD1,2*, P. RAZIN1 & P. HOMEWOOD2,3 1

Institut EGID, Universite´ de Bordeaux, France

2

JVR Centre for Carbonate Studies, Sultan Qaboos University, Oman 3

Present address: 99 bis rue d’Ossau, 64290 Gan, France

*Corresponding author (e-mail: [email protected]) Abstract: The Natih Formation (late Albian–early Turonian, Oman) corresponds to a very broad inner carbonate platform extending over more than 800 km between the Arabian Shield to the south and the Tethys continental margin to the north. Two types of channelized systems have developed recurrently on this inner carbonate platform: † incisions corresponding to strictly erosive drainage systems which formed at the top of the subaerially-exposed platform during significant drops of relative sea level; † tidal channels corresponding to partly erosive but mainly constructive/depositional systems which formed during phases of flooding of the inner platform. The comparative analysis of the basal surface and fill of incisions and tidal channels, based on the study of several outcrops in the Oman Mountains allows the recognition of the similarities and the main differences between these two types of channelized systems which both develop in an inner carbonate platform setting. One of the main criteria of differentiation is the stratigraphic context in which incisions and channels develop. Incisions develop at the top of regressive sequences, whereas the channels analysed here developed during phases of flooding or maximum flooding, during which higher energy processes such as tidal currents developed on the platform. The incision surface is clearly defined, with diagenetic effects such as silicification and dolomitization below, and with the systematic fill of subjacent burrows and cracks by sparitic calcite crystals. The basal erosion surface of channels is often multiple and composite, systematically burrowed, and associated with no significant diagenetic effect. Channels are generally less deep than incisions but their width is very similar. Incisions are longer than channels and present a section that is, on the whole, symmetrical and regular, whereas channels locally have one margin that is steeper and more erosive than the other. Finally, the less diagnostic parameter of differentiation is their fill. Indeed, incisions and channels are similarly filled during phases of flooding of the inner platform. Incisions and channels form significant heterogeneities at the reservoir scale. It is therefore necessary to be able to recognize these two types of channelized systems, in order to predict their geometry, extent and fill type, and the eventual occurrence of associated reservoir bodies in the more distal direction (forced regressive wedges/bioclastic shoals).

The inner part of carbonate platform systems is generally a very flat, shallow-water area which is more or less extensive behind the platform margin. The depositional system in this kind of environment is overall an aggradation of very continuous and low-energy muddy lagoonal facies. At the scale of third-order sequences, the stratigraphic record of this depositional system appears therefore very tabular, suggesting ‘layer-cake’ horizontal correlations of time lines. However, within these third-order sequences, a high level of sedimentary and stratigraphic heterogeneity may

occur, largely linked with fluctuations of relative sea level. The carbonate producers that exist on such shallow water, nearly flat carbonate environments are highly sensitive to slight fluctuations of environmental conditions. A sea-level drop of only a few metres leads to the sub-aerial exposure of large areas on the platform, and a rapid sea-level rise of a few metres may lead to drowning or partial drowning of the platform. In some cases, carbonate production can continue in some areas, whereas inhibition and condensation may occur

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 163–186. DOI: 10.1144/SP329.8 0305-8719/10/$15.00 # The Geological Society of London 2010.

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elsewhere, progressively leading to the creation of intrashelf basins by differential aggradation of carbonate sediments (van Buchem et al. 2002; Gre´laud et al. 2006; Razin et al. 2010). Moreover, on a homogeneously aggrading carbonate platform, a sea level rise of a few metres increases the water depth on the platform and thus the volume of water exchanged by tidal processes. A higher-energy environment is therefore induced, where bioclastic sediments might be laid down within specific sediment bodies. The development and fill of intrashelf basins create complex inclined geometries within the carbonate platform. Erosion and incisions may occur on the carbonate platform and lowstand wedges may develop on its margin during phases of subaerial exposure. High-energy tidal channels and sandwaves may form during phases of flooding of a flat and homogeneously aggrading carbonate platform. All these structures create significant heterogeneities at regional and at reservoir scale which may have considerable impact on stratigraphic correlations and the occurrence of better quality reservoirs.

In this study, the Late Cretaceous Natih Formation of the Sultanate of Oman, Middle East (Fig. 1), has been chosen to illustrate and describe some of these potential heterogeneities within inner carbonate platform systems. The Natih Formation corresponds to a very extensive inner carbonate platform system which developed at the top of a succession of aggrading/prograding carbonate platforms on the Arabian Plate during the Cretaceous (Pratt & Smewing 1993). This formation is a prolific reservoir interval in Oman and in the Middle East and several accessible outcrops of good quality occur in the Oman Mountains (Homewood et al. 2008; Al Jabal al Akhdar, Jabal Madar and the four Jabals of the Adam Foothills, Fig. 1). Abundant subsurface data are also available south and SW of the outcrops to extend the study of this formation and especially to complete regional correlations and understanding. This study focuses on the different channelized systems which developed during the deposition of the Natih Formation and proposes some elements of differentiation between ‘incisions’ and ‘channels’, which developed at different stratigraphic

Fig. 1. Simplified geological map of the central part of the Oman Mountains showing the location of the studied outcrops and the extent of the main correlation transects. (Modified from Gre´laud et al. 2006).

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periods on the Natih platform. Several ‘channelshaped’ sediment bodies are preserved in the deposits of the Natih Formation and their detailed sedimentological and stratigraphic study enables the distinction of two different types, based on the processes which created them. The term ‘incision’ used in this paper (see also Gre´laud et al. 2006) defines a shallow (up to 20 m deep) incised valley, that is ‘a fluvially-eroded, elongate topographic low, typically larger than a single channel form and characterized by an abrupt seaward shift of depositional facies across a regionally mappable sequence boundary at its base’ (Zaitlin et al. 1994). Incisions developed after the deposition of the formation in which they occur (bedrock incision) and are, in the Natih Formation, 1 to 1.5 km wide, 12– 20 m deep, with a symmetrical shape. They form during phases of platform exposure, when the relative sea level drops below the platform margin and are filled during the following sea level rise. The term ‘channel’ used in this paper defines a tidal channel (an erosive/constructive channel formed by tidal processes or currents), contemporaneously with the deposition of the formation in which it occurs. In the Natih Formation, channels are 0.1–1 km wide, 2–10 m deep with a slightly asymmetrical shape. They form during phases of flooding or maximum flooding of the platform and are filled either during relative sea level rise (simultaneously as they form) or during early sea level fall. The stratigraphic record of the development of these two types of channelized systems corresponds to ‘channel shaped’ sediment bodies which may appear identical at first sight, since the fill of incisions and channels can be very similar. The sediment bodies corresponding to incision fills and channel fills are in fact genetically different and should not be misinterpreted, since they are keys for the construction of stratigraphic correlations and for the prediction of the location of associated reservoir bodies. In this paper, the stratigraphy of the Natih Formation is first summarized from the results of previous studies (van Buchem et al. 1996, 2002; Droste & Van Steenwinkel 2004; Gre´laud et al. 2006). Then the two different types of channelized systems which develop during the deposition of the Natih Formation are described. The main characteristics of incisions are reviewed, based on the results of the detailed study of Gre´laud et al. (2006), then the channels are described in detail, on the basis of outcrop observations. Whereas the incisions and incision fills have been described in detail (particularly for Sequence I, e.g. on Jabal Shams) in Gre´laud et al. (2006), the channels have received somewhat less attention. This paper provides the complementary data on the channels but does not repeat all the previous illustrations and descriptions

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of the incisions. Incisions and channels are then compared and the main criteria of differentiation are highlighted. Finally, the implications for reservoir heterogeneity and prediction are discussed.

Stratigraphic setting: the Natih Formation The Natih Formation is the youngest of a succession of very wide shallow-water carbonate platform systems which developed in Oman during the Cretaceous, before the beginning of the Late Cretaceous obduction processes (Murris 1980; Hughes-Clarke 1988). Classically, the Natih Formation is attributed to the late Albian –early Turonian (Simmons & Hart 1987; Smith et al. 1990; Scott, 1990; Kennedy & Simmons 1991; Philip et al. 1995; van Buchem et al. 2002; Homewood et al. 2008) and corresponds laterally to the Mauddud, Mishrif and ShilaifKhatiyah Formations on the Arabian Peninsula (Burchette & Britton 1985; Burchette 1993; Alsharhan 1995; van Buchem et al. 1996, 2002; Terken 1999) and to the Sarvak Formation in Iran (James & Wynd 1965; Alsharhan & Nairn 1988; Sharland et al. 2001; Taati 2005; Razin et al. 2010). The sedimentology and stratigraphy of this formation have been studied extensively in the Jabal Akhdar, Jabal Madar and in the four Jabals of the Adam Foothills (Simmons & Hart 1987; Smith et al. 1990; Philip et al. 1995; van Buchem et al. 1996, 2002; Immenhauser et al. 2000; Schwab et al. 2005; Gre´laud et al. 2006; Homewood et al. 2008) and on subsurface data from the south of the Oman Mountains (Terken 1999; Cortis et al. 2001; Keating 2001; Masaferro et al. 2003; Droste & Van Steenwinkel 2004; Morettini et al. 2005; Gre´laud et al. 2006). This formation corresponds to a very broad, extensive inner platform that developed over more than 800 km between the Arabian Shield in the SW and the Tethys ocean-facing platform margin in the north (Fig. 2) (Murris 1980; Ziegler 2001). It comprises aggrading shallow-marine deposits, dominated by benthic foraminifera and rudists. The Natih Formation corresponds to three fully developed third-order sequences (I, II, III) bounded by a sub-aerial exposure surface at the top (Fig. 3) (van Buchem et al. 1996, 2002). A fourth sequence (IV) has also been defined, but is generally truncated by the major erosion surface at the top of the Natih Formation. Two types of sequences are distinguished within the Natih Formation, based on their different stratigraphic architecture (Fig. 4, van Buchem et al. 1996, 2002; Gre´laud et al. 2006). (1) Sequence I (late Albian –early Cenomanian) and Sequence III (late Cenomanian –early Turonian) record a similar depositional system evolution (Fig. 4a). They commence with a tabular, mixed

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Fig. 2. The depositional setting of the Natih Formation. (a) Simplified palaeogeographic map of the Arabian Peninsula during the deposition of the Natih Formation (late Albian– early Turonian) (modified from Murris 1980 and Gre´laud et al. 2006). See position in Figure 1; (b) Schematic regional cross-section through the Natih carbonate platform which presents a very broad inner platform area extending between the Arabian Shield in the SW and the Tethys ocean margin in the NE. The Natih carbonate platform is overall aggrading in its inner part, except during periods of rapid increase of accommodation during which intrashelf basins develop.

carbonate-argillaceous platform system which grades upward to a very slightly inclined (0.18) muddy carbonate ramp system. Then an intrashelf basin develops by differential aggradation of carbonate sediments during a phase of rapid increase of accommodation on the platform, probably linked with the major transgressive events of the late Albian (Seq. I), and late Cenomanian (Seq. III). As carbonate production exceeds the creation of accommodation, the system starts to prograde and the intrashelf basin is progressively filled by a series of sigmoid sedimentary wedges along which a clear differentiation of facies occurs.

The top of Sequence I is characterized by two regional successive phases of platform exposure of eustatic origin (van Buchem et al. 2002; Gre´laud et al. 2006; Razin et al. 2010). These emersion surfaces are characterized by 12 –17 m deep incisions on the exposed carbonate platform and by the development of forced regressive wedges of limited lateral extent on the margins of the intrashelf basin (Gre´laud 2005; Gre´laud et al. 2006). In contrast, The top of the Natih Formation is characterized by a phase of platform exposure of tectonic origin. It is due to active uplift of the northeastern Oman margin through the flexure of the

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Fig. 3. Table summarizing the stratigraphic location and main characteristics of the different channelized sediment bodies which develop within the Natih Formation. According to their different parameters and mainly to their stratigraphic position, two types of channelized systems are differentiated. Incisions (i) develop during periods of sub-aerial exposure of the carbonate platform, whereas tidal channels (c) develop during periods of flooding of the platform. The section on the left is a composite outcrop section.

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Fig. 4. Stratigraphic transects through (a) Sequence I (modified from Gre´laud et al. 2006) and (b) Sequence II, based on the analysis and correlation of outcrops in the east (Adam Foothills) and subsurface data in the west (well logs, cores, and seismic surveys). These correlation transects illustrate the two different types of depositional system in the Natih Formation: (1) a differential aggradation system leading to the creation and fill of an intrashelf basin in sequence I (similar in sequence III); and (2) a homogeneous aggradation system leading to the deposition of very continuous carbonate beds made of shallow-marine/lagoonal facies in Sequence II. Channelized systems develop within both depositional systems: incisions occur at two stratigraphic levels in the upper part of sequence I (IS1 and IS2), and channels develop within the fill of the incisions, and at the base as well as in the upper part of sequence II (within sequence II-1 in Jabal Madar and below the top of Sequence II-2 in Jabal Madmar). Note the very high vertical exaggeration on these transects.

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Arabian plate at the end of the Cenomanian and early Turonian (Patton & O’Connor 1988; Scott 1990; Burchette 1993). This emersion surface is characterized by deep incised valleys (up to 180 m), identified both at outcrop (Jebel Madar, Jebel Akhdar; van Buchem et al. 2002) and on seismic data (Droste & Van Steenwinkel 2004). They are filled with terrigenous or estuarine to lagoonal carbonate deposits. (2) Sequence II (middle Cenomanian) also starts with a very tabular mixed carbonate-argillaceous platform system with a progressive decrease of clays upwards, but no intrashelf basin develops during the transgressive phase of this sequence, probably due to a slower increase of accommodation on the platform at this time (Figs 3 & 4b). The depositional profile remains flat during this sequence, although low amplitude (10 to 20 m) and large (more than 50 km) wavelength tectonic deformations have been inferred on the basis of thickness variations across the studied area. This sequence is characterized at the base by a succession of high frequency cycles made of relatively thin (1–2 m thick), very extensive (over more than 100 km) and very tabular carbonate beds, mostly made of muddy carbonate facies, deposited under low-energy, shallow-marine conditions. They correspond to the aggradation of the platform during high-frequency phases of increase of accommodation. In the upper part of the sequence, the carbonate beds become thicker and are made of more open (higher water depth, higher energy, cleaner conditions) lagoonal facies with larger organisms such as in-situ rudists, stromatoporoids and corals. These observations show that more accommodation was available on the platform in the upper part of the sequence. At the same stratigraphic level, interpreted as the maximum flooding level, high-energy bioclastic channels, sandwaves, and larger shoals occur. Sequence II is therefore characterized by asymmetrical accommodation cycles, since the sediments are mainly deposited by aggradation during the transgressive part of the cycles. There is an emersion surface at the top of this sequence, characterized locally by small root traces and pedogenesis.

Incisions In the Natih Formation, incisions occur at two stratigraphic levels in the upper part of the first third-order sequence (top Sequence I-6 and top Sequence I), and at the top of the Natih Formation (Fig. 3). The incisions at the top of the Natih Formation are large and deep incised valleys developed as a

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result of a flexure associated with the thrusting of the Oman nappes. These incisions range from several hundreds of metres to more than fifty kilometres wide and from tens of metres to more than one hundred metres deep. They are observed on 3D seismic data (Droste & Van Steenwinkel 2004) and interpreted from outcrop correlations (major truncation east of the Jabal Akhdar and in Jabal Madar area; van Buchem et al. 2002). They are partly filled with iron-rich sandstone in Jabal Madar. These incisions are regionally overlain by the Santonian deep-water Muti marls which mark a widespread regional transgression (Robertson 1987). Locally, iron-oolitic deposits as well as carbonate bioclastic channels are found in the incisionfill, below the Muti marls (e.g. Wadi Mi’Aidin, van Buchem et al. 2002). The upper-Sequence I incisions are described in detail in Gre´laud et al. (2006) on the basis of the analysis of several outcrop examples (Jabal Madar, Madmar and Shams, Fig. 1) and the interpretation of 3D seismic reflection profiles and amplitude maps. These incisions occur along fourth- and third-order sequence boundaries, at the top of shallowing-upward sedimentary cycles. They develop during phases of sub-aerial exposure of the platform and are associated with forced regressive wedges which prograde on the margins of the intrashelf basin (Fig. 4a).

Main characteristics of incisions Geometry. The incisions in the upper part of Sequence I are 1 km wide, between 10 –20 m deep, with a fairly regular and symmetrical crosssection (Figs 3 & 5). They appear as single streams with no tributaries and extend over at least 40 km to more than 100 km on the originally flat carbonate platform top. Their morphology is slightly sinuous to sinuous with irregular incised meanders which are controlled locally by structural features such as faults or salt diapirs, passing around the positive elements (based on both outcrop and seismic data; Gre´laud et al. 2006). Basal surface. The erosion surface at the base of the incisions is single and sharp (Fig. 5d). It is associated with intense silicification along, below and/or above the surface with frequent large chert nodules (the cause of silicification is currently unknown and under study). Subtle petrographic features indicative of emersion such as micro-karst, sparite-filled burrows and cracks, microbreccia and dolomitization are locally seen on or below the incision surface. Incision fill. The incisions are filled during the transgressive phase which follows the period of

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Fig. 5. Review of the main characteristics of Natih Sequence I incisions (modified from Gre´laud et al. 2006). (a) Schematic correlation of the stratigraphic unit between the two incision surfaces in the upper part of Sequence I (Sequence I-7) between Jabal Madar and Jabal Madmar. This diagram illustrates the fill of IS1 incisions on a proximal-distal transect. The incisions are filled by three successive facies units: (1) a high energy lag deposit at the base, followed by (2) a low-energy muddy facies above and topped by (3) a higher-energy facies unit corresponding to tidal channel deposits. (b) Schematic interpretation of the Natih incision fill succession. The lag was probably deposited during a phase of tidal erosion at the first flooding of the platform after exposure; the muddy unit may correspond to a phase of development and migration of bioclastic shoals on the platform margin which would protect the incisions from

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platform exposure. The incision fill comprises three main facies units the boundaries of which are characterized by progressive evolution of facies (Fig. 5a). The first unit at the base is a high-energy lag deposit made of platform weathering products preserved as carbonate grainstone/rudstone to microconglomerate with very rare quartz grains. The second unit corresponds to low-energy confined deposits, either made of mudstone or green clays containing sepiolite (chemically-precipitated magnesium-rich clay mineral). This second muddy unit is often partially to entirely dolomitized in the field. The third fill-unit is made of less-restricted and higher energy deposits, normally organized as large low-angle inclined stratifications (Fig. 5c). They are interpreted as laterally-accreting point-bar deposits within shallow tidal channels which developed when the whole platform was flooded (Fig. 5b). These tidal channels re-occupy the remnant depression at the top of the incision fill. This last unit is capped by a more or less well developed hardground. Considering the thickness of their fill (i.e. depth of erosion, up to 20 m) and the strong downward stepping pattern of the forced-regressive wedges, these incisions appear to be the result of 20–30 m eustatic sea-level falls. They are filled during the following rise of sea level with intertidal to subtidal facies, dominated either by carbonates when the carbonate factory resumed directly at the beginning of the transgression, or by clays when the carbonate factory was inhibited after a probable longer phase of platform exposure which might be linked with a higher amplitude of the previous sea-level fall.

Channels In the Natih Formation, bioclastic-grain filled channels are observed in three different settings (Fig. 3): (1) at the top of the fill of upper Sequence I and top Natih incisions, (2) in the upper part of highfrequency cycles within Sequence II-1, and (3) associated with bioclastic sandwaves in the upper part of Sequence II (Fig. 4b). They all develop during a phase of increasing accommodation on the platform and they all have an aggrading component.

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Channels at the top of incision-fill Bioclastic channels develop within the fill of upper Sequence I and top Natih incisions. These channels are located at the base of the transgressive systems tract above major emersion surfaces. Fill of upper Sequence I incisions. The study of the outcrops shows that in some cases channels develop during the final fill of the incisions (third fill-unit of IS1 in Jabal Madar and Jabal Madmar, third fill-unit of IS2 in Jabal Shams). The base of these channels is either a simple and single erosion surface corresponding to a tidal ravinement surface (base of third fill-unit of IS1 in Jabal Madar, Fig. 5a), or a succession of several poorly marked, composite surfaces, suggesting a clear aggradational trend (base of third fill-unit of IS1 in Jabal Madmar, Fig. 5a and base of third fill-unit of IS2 in Jabal Shams, Gre´laud et al. 2006). These channels cut through bioturbated mudstone with orbitolinids deposited during the second phase of incision infill. They develop over the whole width of the re-occupied incision, that is, over 1 km, in all of the outcrops studied. In Jabal Madmar and Jabal Shams, the channelized last unit of the incision fill is a 3 m-thick interval with laterally-accreting strata (Fig. 5c). Each inclined bed shows a lateral facies evolution from floatstone with large organisms (broken or in situ) at the top (rudists, stromatoporoids, bivalves and small corals) to wackestone with rudist fragments and finally to marly mudstone at the toe of the beds. These low-angle inclined stratifications are interpreted as low-energy ‘point bar’ deposits within shallow tidal channels which are reoccupying the incisions during the flooding of the whole platform top (Figs 3 & 5b). In Jabal Madar, the base of the channelized last unit of the incision fill is a composite erosion surface, cutting down 10 m into the underlying transgressive muddy deposits (Fig. 5a). Above this erosion surface, several successive metre-thick, oblique (few degrees) coarse grainstone units are vertically and laterally stacked towards the SE. They comprise subrounded lithoclasts, shallow benthic foraminifera and bivalve and echinoderm fragments with mega-ripple cross-bedding. On the margins of this erosive channel unit, the grainstone beds thin out and grade laterally to a 3 m thick

Fig. 5. (Continued) high-energy conditions; the development of tidal channels re-occupying the incisions in the upper part of the fill probably corresponds to the phase of flooding of the platform, during which a larger volume of water was available to amplify or renew tidal processes on the platform. (c) The tidal channel deposits in the upper part of the incision fill are characterized, in the distal position (Jabal Madmar), by lateral accretion geometries, interpreted as the result of low-energy ‘point bar’ migration. (d) Incisions are characterized by a very sharp and single basal erosion surface, which is often found silicified in the field.

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interval of highly bioturbated wackestonepackstone containing abundant shallow benthic foraminifera and in situ rudists, pectens and stromatoporoids. The sediments within these channelized units are interpreted as high-energy deposits in a shallowmarine channel. No direct diagnostic evidence suggesting tidal processes has been observed but cross-bedding indicates a fluctuating unidirectional current (east –west) that is most probably of tidal origin. The channelized erosion surface is therefore interpreted as the result of multiple tidal ravinement surfaces that developed during the maximum flooding stage in the lagoon. The channels seem to re-occupy a possible remnant depression at the top of the incision fill, where tidal current energy would concentrate. On the margins of the active channel, transgressive muddy sediments aggrade in a shallow-lagoon environment. Fill of incisions at the top of the Natih Formation. The large and deep incisions at the top of the Natih Formation have been shown on the southern flank of the Jabal Akhdar by stratigraphic correlation of several field sections (van Buchem et al. 2002, p. 38, Fig. 9). The fill of these incisions corresponds there to a complex assemblage of carbonate units that are more or less channelized, but the geometry of which is complicated by syn-sedimentary tectonic deformation. Bioclastic channels can be observed at the outcrop scale, as for example in Wadi Mi’Aidin (Fig. 6). In Wadi Mi’Aidin, the main incision surface at the top of the Natih Formation is marked by an iron oxide crust and cuts into the facies with rudists of the third Natih sequence (Seq. III). The incision appears to be relatively deep since several stratigraphic units are missing (like Sequence IV,

van Buchem et al. 2002). The incision fill starts with the aggradation of nodular and argillaceous carbonate deposits of shallow-water facies. These deposits are cut by the base of a channel which is several hundreds of metres wide and five metres thick (maximal visible thickness). The fill of the channel is made of bioclastic cross-bedded grainstone. The top of this incision-fill unit is marked by a 1 m thick layer of ferruginous oolites, directly overlain by deep-water hemipelagic marls of the Muti Formation. Similar bioclastic channels are also visible at the same stratigraphic level to the west of the Jabal Akhdar on Wadi Khamah, Wadi Sumayt and Wadi Tanuf sections (van Buchem et al. 2002). Stratigraphic signification of the channels in the incisions. Given that incisions result from an emersion of the platform in response to a drop of relative sea level, the channels which develop within the fill of these incisions belongs to the transgressive systems tract at the base of the following sequence. They therefore developed during a phase of increased water level on the inner platform and record an increase of energy in the environment (Fig. 5b). Channels with an erosive base are interpreted to develop during phases of maximum flooding, when the maximal volume of water was driven by tidal currents in this proximal setting. Channels with a less clear, composite base develop preferentially during the transgressive phases during which the energy level in the system was lower but the aggrading trend already relatively high (Fig. 5a).

Bioclastic channels of Sequence II-1 The studied outcrop, located on the southern flank of Jabal Madar, allows the observation of a bioclastic

Fig. 6. Channel geometries observed at the top of the Natih Formation in Wadi Mi’Aidin (SE of the Jabal Akhdar). The bioclastic channel eroding the upper part of Sequence III develops within the fill of a large-scale incision at the top of the Natih Formation.

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channel at the top of the second high-frequency cycle of Sequence II-1 (Fig. 3). The first highfrequency sequence of Sequence II-1 corresponds to the fill of top Sequence I incisions. The second high-frequency sequence of Sequence II-1 is recognized throughout the whole Foothills area (Jabal Qusaibah to Jabal Madar). It starts with lagoonal green shales located just above the hardground marking the top of Sequence I (Fig. 7). These shales grade progressively upwards to highly nodular argillaceous mudstone with orbitolinids and then to a 4 m thick carbonate bed made of nodular wackestone with gastropods and lamellibranches, and numerous large rudists (radiolitids, followed by caprinids) in life position. The evolution of rudist associations indicates an increasingly open platform setting. This carbonate bed is capped by a hardground of regional extent which corresponds to the top of this second high-frequency sequence. This sequence records the transition from argillaceous sedimentation in a restricted marine environment to carbonate sedimentation in an increasingly open inner platform setting in a transgressive context. This sequence is asymmetric since the regressive phase of the cycle is almost only recorded by the hardground surface. At the Jabl Madar outcrop location, a bioclastic channel develops in the upper part of this sequence (Fig. 7). On the western margin of the channel, the initiation of the erosion surface starts a few decimetres below the hardground and cuts through the whole sequence down to peritidal cycles that are located in the upper part of Sequence I (Sequence I-7). This channel is approximately 750 m wide with a maximum depth of about 10 m. It shows an asymmetric cross-section with a western margin steeper than the eastern margin. No orientation can be given for this channel since it is only a 2D exposure. The fill of this channel comprises a complex set of channelized bodies, mainly of fine –coarse grainstone with rudist fragments, other bivalves, corals and echinoids. These bioclastic facies generally show planar horizontal or trough cross-bedded stratifications indicating a main current direction towards the NE. The sigmoid shape of the cross beds (mega-ripples) and the (scarce) occurrence of mud drapes on some sets of laminae, suggest a tidal influence. The top of this bioclastic fill corresponds to an interval of yellow clays, locally overlain by a conglomerate with lithoclasts and large fragments of corals and rudists. Dolomitized vertical burrows are visible in the finer-grained intervals and at the base of the channelized bodies. In detail, two small (3 m thick) units with a similar vertical evolution of facies can be identified in the channel fill. These units are arranged by

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vertical aggradation in the channel. They start with an erosion surface generally overlain by a coarse lag, followed by fine grainstone deposits with horizontal stratification, grading progressively, or with an erosive contact, to coarser deposits with trough cross-bedding and lens-shape stratification cut by multiple erosion surfaces. Therefore, each of these small ‘very high frequency’ sequences clearly records an increase of current velocity in the channel. On the eastern margin of this channel, the basal erosion surface becomes less and less sharp until a lateral variation of facies occurs between the grainstone at the top of the channel fill and the wackestone with rudists of the substratum bed (Fig. 7). This lateral change of facies shows that the channel developed simultaneously with the upper part of the second high-frequency sequence of Sequence II-1. Several much smaller channels (100 m wide, 0.7 m thick) are found at the same stratigraphic level to the west of the main channel from which they are separated by muddy rudist bank facies. They are made of medium to coarse grained grainstone with cross-beds and large fragments of Icthyosarcholithes rudists. Interpretation. The succession of facies in the second high-frequency sequence of Sequence II-1 records a transgressive trend during which the sedimentary system evolved from a restricted marine environment with confined argillaceous sedimentation to an increasingly open inner carbonate platform environment. The studied channel developed within the upper part of this sequence of facies. This sediment body which incises the shallow and muddy rudist banks testifies that powerful tidal currents developed on the inner platform. The onset of these currents appears to have been directly linked with the flooding of the platform and therefore with an increase of the volume of water in movement in the system. Unlike the incisions, this erosive structure corresponds to a rise of relative sea level. In detail, the fill of this channel suggests a multiphase evolution. Several flooding phases associated with an increase of current energy are recorded. The basal erosion surface of the channel is therefore probably slightly diachronous and formed progressively while the platform aggraded (constructive/ depositional channel).

Bioclastic channels and sandwave complexes at the top of Sequence II Jabal Madar. In Jabal Madar, the top of Sequence II is always underlined by a highly dolomitized 1 m thick interval. Detailed correlations at the 10 km

174 C. GRE´LAUD ET AL. Fig. 7. Outcrop correlation transect showing the geometry and fill succession of a tidal channel located at the base of Sequence II (base of sequence II-1) and cropping out on the southern flank of the Jabal Madar. Note the highly asymmetric accommodation cycle within this high-frequency sequence at the base of the third-order transgressive systems tract of Sequence II.

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scale of the Jabal show that this limit corresponds to a truncation, since a 2 m thick bed has disappeared through erosion between the eastern and the western flanks of the Jabal. Very locally, a 20 m wide and 70 cm thick channel of fine dolomitized grainstone crops out just below this surface. It incises shallow-water inner platform facies, represented by bioturbated wackstone with praealveolinids and small solitary corals. Jabal Madmar. In the northwestern part of Jabal Madmar (Madmar 6 wadi), a complex of bioclastic channels develops in the upper part of Sequence II-2 (Fig. 8). It is formed by the eastward offlapping lateral stack of lens-shaped channelized bodies, 0.5–1 km wide and up to 4 m thick (Fig. 9). This group of channels is localized over a limited area, 2–3 km wide and oriented east –west. These channels develop within a higher-energy lagoonal facies containing in situ rudists, corals and stromatoporoids, forming the last bed of Sequence II. These channels are filled with coarse bioclastic grainstone with dominant rudist fragments, which can grade laterally and vertically to fine

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grainstone-packstone with silicified nodules (Figs 9 & 10b). The grainstone facies shows planar crossbedding structures, often poorly visible at the outcrop. The geometry of the erosion surface and the relations with the underlying facies of the substratum are comparable to those described previously for the base Sequence II-1 channel. On the margins of the channel complex, the basal surface is slightly inclined. The surface is either clearly erosive (western channel margin, Fig. 8c) or characterized by multiple minor erosion surfaces and lateral variations between the bioclastic channel-fill facies and the adjacent muddy rudist-bank facies (eastern margin, Fig. 8b). This surface is highly burrowed and the burrows penetrating in the subjacent unit are filled by the bioclastic grains coming from the channel fill (Fig. 10a). This observation shows that the subjacent deposits were not completely lithified during the development of the channel and confirms the absence of any significant discontinuity between the channel and the surrounding substratum facies, in contrast with the incisions. In detail, the architecture of this channel system is relatively complex (Fig. 9). On the western

Fig. 8. (a) Outcrop view of the bioclastic channel cropping out in the NW of the Jabal Madmar (Madmar 6 wadi) just below the top of Sequence II. (b) Eastern margin of the channel where laterally accreting geometries are observed at the same stratigraphic level as the channel fill unit. (c) Western margin of the channel where clear erosion of the underlying strata can be seen.

176 C. GRE´LAUD ET AL. Fig. 9. Outcrop correlation transect showing the detailed geometry of the upper-Sequence II channel in Jabal Madmar. A clear erosive margin is observed on the western margin of the channel, whereas a progressive evolution of facies from aggrading lagoonal deposits to coarse bioclastic channel-fill is seen on the eastern margin of the channel.

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Fig. 10. (a) Highly burrowed and irregular basal surface of the upper-sequence II channel in Jabal Madmar. These burrows suggest that the substratum was not completely lithified at the time of development of the channel. (b) Progressive vertical evolution of facies from a thin wedging grainstone bed at the base of the channel fill to a dolomitized packstone/grainstone unit on the western margin of the channel. This latter nodular unit is eroded towards the east by the main channel, later filled with bioclastic carbonates.

margin, a first sigmoid body, made of coarse grainstone drapes the erosion surface. This bioclastic facies grades laterally and vertically upwards to an interval with a more muddy bioturbated, dolomitized and silicified facies (Fig. 10b). Towards the east, this facies is cut by the basal erosive surface of a second channel with a very coarse bioclastic fill (Fig. 9). The eastern margin of this channel is not sharp or single. It is characterized by a lateral facies transition between these coarse granular deposits and the muddy deposits with rudists. This transition takes place along broad low-angle inclined stratifications suggesting a process of lateral accretion on the margin of the channel (Fig. 8b).

channels, as shown by the lateral accretion processes. These characteristics are very different from those described for the incisions. However, the nature and geometry of the channel-fill deposits show some analogies with those of the fill of the incisions. Indeed, a first grainstone deposit grading upward to a dolomitized mudstone, itself cut by a grainstone channel with cross-bedding, also forms the fill of the first incision in Jabal Madar (Fig. 5a, Gre´laud et al. 2006). These analogies show that the channels and the incisions were both filled under a comparable context: a transgressive phase allowing the flooding of the inner platform and the onset of high-energy processes.

Interpretation. As with the previous example, the occurrence of this channel system in the upper part of Sequence II shows the development of strong currents during the maximum flooding period of the inner platform (no current directions could be measured with precision). The tidal origin of these currents could not be established here. However, the relationships with the substratum deposits show that at this stage, the muddy rudist and stromatoporoid banks were drained and eroded by relatively strong currents, concentrated in these more or less erosive and probably sinuous

Jabal Qusaibah. In Jabal Qusaibah, a 3– 4.5 m thick bioclastic sandwave complex (apparently migrating towards the NE) is found exactly at the same stratigraphic level as the channels in Jabal Madmar, in the topmost part of the second third-order sequence of the Natih Formation (Fig. 11). This Sequence II, from which only the upper part crops out (Seq. II-2), shows similar facies and similar stratigraphic units in Jabal Qusaibah as in Jabal Madmar, located 50 km towards the east (Fig. 12a). In Jabal Qusaibah, the thickness of the beds is slightly greater, and the nature and geometry

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Fig. 11. Outcrop part of the Sequence II correlation transect of Figure 4b highlighting the similar stratigraphic location of the Jabal Madmar channel and the Jabal Qusaibah sandwaves: at the maximum flooding of Sequence II. Asymmetrical accommodation cycles are also observed at the scale of fourth-order sequences.

Fig. 12. Outcrop views of the sandwave complex located on the northern flank of the Jabal Qusaibah. (a) Overview of the stratigraphy of the Jabal Qusaibah outcrop. (b) Photograph showing the sandwave complex and its northeastern termination. (1) lower unit, (2) middle unit, (3) upper unit (see Fig. 11) (c) This complex of coarse bioclastic sandwaves is up to 5 m high and presents steep foresets migrating in various directions (overall towards the NE).

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of facies only vary in the topmost part of the sequence (Fig. 11). The equivalent of the 3 m thick higher-energy lagoonal unit, where the channels of Jabal Madmar develop, corresponds here to a 9 m thick interval comprising three units on the reference section: (1) a lower unit of packstone with prae-alveolinids, stromatoporoids and rudist fragments; (2) a middle unit with a sharp contact on the previous one, composed of very fine grainstone-packstone with colonial corals, stromatoporoids and rudists (Ichtyosarcholithes); and (3) an upper unit made of very nodular wackestone with benthic foraminifera and a dolomitized top marked by very small root traces (top Sequence II, van Buchem et al. 2002). Laterally towards the SW, the middle unit is made of bioclastic sandwaves, several metres thick (3–4.5 m) similar in composition to the contemporaneous channels of Jabal Madmar: grainstonerudstone with rudist and chondrodont fragments (Fig. 12). These sandwaves are built up by of large inclined foresets (20– 258) sloping from the top to the base of the complex (class I dune of Ashley 1990). However, the sandwaves are composite since they are formed by the amalgamation of several bioclastic bodies with large foresets of different and sometimes opposite orientations (e.g. class V sandwaves of Allen, 1980). The front part of this dune was fossilized by the fine grainstone deposits described on the reference section (unit 2): they onlap the uppermost (youngest) foreset (Fig. 12). Interpretation. The depositional geometry and the multidirectional character of the currents suggest that these structures were built by strong tidal currents. The height of these bioclastic sandwaves indicates a depth of the order of 15 m. Such a bathymetry confirms that these bioclastic sandwaves developed during a major phase of flooding of the platform. The succession of facies within Sequence II shows that this stratigraphic interval represents the period of maximum flooding (Fig. 11). The upper continuous unit (3) of nodular wackestone with foraminifera shows a decrease of bathymetry and corresponds to the highstand systems tract of Sequence II-2. The regressive trend of this unit becomes much more pronounced at the top, since the presence of root traces indicates sub-aerial exposure of the platform. The occurrence of channels 50 km towards the east in Jabal Madmar, at the same stratigraphic level as the bioclastic sandwaves, shows that the depth was probably less in Jabal Madmar area, which fits in well with the east–west trend of the system established for the whole Natih Formation on the Foothills transect (van Buchem et al. 2002). The dip of the depositional profile on the platform

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stays extremely low during this period since the bathymetry only increases of a maximum of about 10 m over 50 km (0.018, Fig. 11). However, all these correlations confirm again that the channel system of Jabal Madmar develops during a phase of major flooding of the inner platform. Jabal Shams. Compared to the Adam Foothills, the Jabal Shams location corresponds to an area that was closer to the ocean margin of the platform (Figs 1 & 2). Sequence II-2, which is reduced in thickness to only 13 m, has been analysed on several sections in the Wadi Nakhr canyon and on the Jabal Shams plateau (Fig. 13). On the reference section called the ‘canyon section’, this sequence is made up of a 4.5 m thick unit (1) of two nodular bioturbated wackestone beds with stromatoporoids, separated by a thin argillaceous mudstone layer. The top of this unit is marked by a highly bioturbated surface, interpreted as a ‘firmground’. It is overlain by a second unit (2), 2.5 m thick, starting with a dolomitized mudstone layer (d) (0.5 m) overlain by a nodular bioturbated mudstone bed. Above, a 7 m thick bed is found (3), the composition of which evolves rapidly from a slightly argillaceous bioturbated mudstone, dolomitized at the base, to a coarsening-upwards bioclastic grainstone, and finally to a floatstone with large rudists (m). A thin bed of fine grainstone (30 cm) capped by a regionally extensive perforated hardground marks the top of Sequence II. On the Jabal Shams plateau, 3 km SW of the ‘canyon section’, a system of bioclastic bedsets developed above the first unit of wackestone with stromatoporoids (within unit 2). These bedsets form a 10 m thick, 600 m long sediment body, made of very coarse grainstone and rudstone with rounded rudist fragments (Fig. 13). These highenergy deposits are characterized by large inclined stratifications with amplitudes of several metres and dip up to 208 towards the NE. This sediment body includes a vertical and lateral succession, from the SW towards the NE, of several coarse bioclastic sigmoidal beds separated by finer mudstone-wackestone intervals, which are bioturbated and some are slightly dolomitized. These bioclastic sandwaves start in the SW with a lower unit (a) made of very fine bioturbated and dolomitized mudstone facies, which clearly thin out towards the NE, building an initial inclined palaeotopography. This geometry may be the result of a differential aggradation of the sediments on the platform linked with a transgressive phase, or may correspond to the toes of foresets from another bioclastic shoal which cannot be observed due to outcrop conditions. Three coarse bioclastic bedsets of sigmoid shape (b1 to b3) are then laterally stacked towards the NE. They are large foresets

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Fig. 13. Natih sequence II bioclastic sandwave complex of Jabal Shams: geometries and facies. (a) Outcrop correlation transect. (b) Large sigmoid foresets building the sandwave complex.

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within bioclastic sandwaves, made of very coarse rudstone with centimetre-wide sub-rounded rudist fragments grading laterally and downwards to fine bioturbated and dolomitized mudstone/wackestone (bottomsets). The base of the foresets is locally erosive. Multidirectional 3D cross-sets are well expressed in the grainstone facies along the foresets. These three sigmoidal bedsets (b1 to b3) grade towards the NE into a fine and dolomitized bottomset facies, which is equivalent to the dolomitized layer of the reference section (d). Finally, a more tabular, sedimentary body of rudstone and/or grainstone (b4) compensates for the remnant topography at the top of this sandwave complex (Fig. 13a). It is overlain by a metre-thick nodular wackestone bed with prae-alveolinids (w), itself overlain by a 1 m thick coarse grainstone bed (b5), topped by a thin marker horizon (m) made of floatstone with large and whole rudist shells. This marks the top of the sequence. Interpretation. Although clearly thinner, Sequence II-2 in Jabal Shams globally shows an evolution of facies similar to the one described in the Adam Foothills. In fact, this sequence starts here with muddy facies containing stromatoporoids, typically corresponding to a low-energy lagoon environment and ends up with coarse bioclastic facies including large rudists showing an increase of energy. This increase of energy locally leads to the development of bioclastic sandwaves on the platform, the amplitude of which shows that the water level was relatively high on the inner platform. The onset and migration of these sandwaves are therefore interpreted as the record of a maximum flooding phase on the platform, at the scale of Sequence II. Subsurface. The correlation of outcrop data with subsurface data west of the outcrops (well logs and cores) shows that well-developed bioclastic shoals occurred at the same stratigraphic level as channels and sandwaves in the upper part of Sequence II (Fig. 4b). These bioclastic shoals are up to 12 m thick and may have been connected. They occur in areas where more subsidence was available during the deposition of the upper part of Sequence II-2 compared to areas where the channels and isolated sandwaves occur. The carbonate grainstone shoals grade laterally towards the east, that is, towards areas where there was less subsidence, to higher-energy lagoonal facies containing in-situ rudists, corals and stromatoporoids. The Jabal Madmar channel and the Jabal Qusaibah sandwaves develop in this higher-energy lagoonal facies.

Main characteristics of channels Geometry. The channels within the Natih Formation are 0.5–1 km wide and are several decimetres to

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10 m deep, but more commonly 2–5 m deep (Figs 3 & 14). They generally crop out with an asymmetric cross-section, having one margin steeper than the other. They are meandering channels and appear on subsurface seismic amplitude maps as highly sinuous. They cannot be correlated between outcrops, suggesting that they have a limited extent (less than a few tens of kilometres). It is currently not possible to define precisely in which direction these channels were connected to the open sea (probably towards the north or the east?). Basal surface. The basal surface of the channels is either clearly erosive but less and less sharp towards the margins of the channel where multiple stacked erosion surfaces are visible, or not clearly marked and corresponding to the toes of inclined laterally-accreting bedsets (low-energy ‘point bars’, particularly the last fill unit of incisions). Where the basal surface is erosive and visible, it has been burrowed and the burrows are filled with the overlying bioclastic channel-fill facies. This observation suggests that the substratum was not lithified at the time of development of the channels. No significant diagenetic modification associated with channel basal surfaces is observed. Channel fill. The studied channels are filled with high-energy facies, rudstone, grainstone or packstone, alternating sometimes with finer and often dolomitized facies. The channel fill is organized as a complex set of carbonate bioclastic sediment bodies, generally corresponding to laterallyaccreting bedsets with low-angle inclined stratifications. A gradual variation of facies between the bioclastic channel fill and the interchannel muddy deposits is frequently observed, at least on one margin of the channels. This observation suggests that the channels developed at the same time as the lagoonal muddy sediments aggraded on the inner platform. The channels studied in the Natih Formation are therefore interpreted as having formed during transgressive phases by tidal ravinement and accretion in tidal creek depositional systems.

Differentiation between incisions and channels The main criterion of differentiation between incisions and channels is the stratigraphic context in which they develop (Fig. 14). † Incisions develop at the top of regressive sequences, generally recording a decrease of energy in the depositional environment and a decrease of carbonate production. At the regional scale, incision surfaces are contemporaneous with

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Fig. 14. Table summarizing the differences between the two types of channelized systems that developed on the Natih carbonate platform. (Modified from Gre´laud et al. 2006.)

forced regressive wedges on the margins of the intrashelf basin. † Channels develop during flooding of the inner platform. They therefore occur at the top of a transgressive sequence of facies, which records an increase of the rate of creation of accommodation on the platform, an increase of the energy level of the depositional processes and commonly an increase of carbonate production. At the regional scale, channels are stratigraphically associated with the development of bioclastic sandwaves which develop in the areas of the inner platform where more accommodation space is available. This difference of stratigraphic context at the time of development of incisions and channels has an influence on several parameters and especially on the nature of their basal erosion surface and their associated structures (Fig. 14).

† The erosion surface at the base of incisions is clearly defined, marked by diagenetic processes such as silicification and dolomitization, and by the systematic fill by sparry calcite crystals of burrows and cracks in the bedrock below. † The erosion surface at the base of channels is commonly multiple and systematically bioturbated. No significant diagenetic modification is associated with this surface. The erosion surface is generally less sharp on the margins of the channels, and particularly on the less steep margin. A gradual variation of facies is observed along this margin between the bioclastic facies in the fill of the channel and the more muddy facies of the surrounding rocks. The morphological differences between incisions and channels are more subtle. Channels are generally less deep than incisions, but they can reach depths comparable with those of incisions (e.g.

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Fig. 15. Impact of various types of channelized bodies on reservoir architecture and heterogeneities.

10 m for the channel at the base of Sequence II-1 in J. Madar). Their width is very similar, of the order of 1 km. † Incisions have an overall symmetrical and regular cross-section, whereas channels may have one margin steeper and more erosive than the other. † The incisions observed on seismic amplitude maps have a slightly sinuous morphology. Interpretation of highly sinuous channels on Sequence I seismic data have been observed (Vahrenkamp et al. 2006) but no images have been published to date. Small highly sinuous channels have been interpreted on amplitude maps made from the top Natih seismic reflection (Droste & Van Steenwinkel 2004, p. 199, Fig. 14). They might correspond to the channels observed in the outcrops. The sediment bodies deposited as large oblique or sigmoid stratifications within the fill of the channels indeed suggest the presence of meanders and laterally accreting point bars. † Incisions have a longer extent than channels, since they can be correlated over more than 100 km. Channels are more difficult to correlate between the different outcrops and have an extent generally less than a few tens of kilometres. Finally, the least relevant parameter in differentiating incisions from channels is their fill. Both of these channelized systems were filled during phases of flooding of the inner platform.

Implications for reservoir heterogeneity and prediction Incision-fills form sinuous elongated sediment bodies which are 1 km wide, 12 –20 m thick and 40 to .80 km long. Channels form sinuous to highly sinuous elongate sediment bodies which are 1 km wide, 3–10 m thick and 20–30 km long. Incisions are relatively frequent within specific stratigraphic intervals (upper Sequence I, i.e. upper Natih-e member and top Natih, Fig. 3) since they have been observed in all the outcrops studied and at several locations on subsurface seismic amplitude

maps. Channels are frequent in the last fill-unit of incisions and at the base and top of transgressive cycles (high frequency to third-order cycles) within the second sequence of the Natih Formation. Incisions and channels have a strong potential for creating a high level of horizontal heterogeneity in the stratigraphic interval in which they occur (Fig. 15). They may lead to a compartmentalization of the reservoir when they are filled with nonreservoir facies like shales. They may form isolated reservoir units when they are filled with reservoir facies and occur within non-reservoir stratigraphic intervals. They may also form permeability streaks when filled, or partly filled with high-permeability reservoir facies.

Incisions Reservoir properties and creation of heterogeneity. Incisions are filled with three successive and different fill units (Figs 5 & 14). The first part of the fill corresponds to a good reservoir facies and may form a high-permeability streak in a reservoir interval. The second and main part of the fill corresponds to poor to non-reservoir facies and may lead to a compartmentalization of part of the reservoir. The last part of the fill corresponds to higher-energy and better reservoir facies, but this unit can be complex incorporating both stacked reservoir and non-reservoir facies. Diagenetic alterations can modify the original properties of incision-fill facies. In the field, the mudstone interval in the second fill unit is commonly dolomitized or partially dolomitized, and crops out as sucrosic dolomite which presents very good potential reservoir properties. The basal surface of incisions is highly silicified in the field and may form a permeability barrier or baffle and disconnect the incision fill from the surrounding rocks. Prediction. The identification of incisions is very useful to detect sequence boundaries and interpret their stratigraphic order (Gre´laud et al. 2006). The presence of incisions can also be used to predict the existence of high-angle forced regressive wedges which would be located on the margin of

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the platform, in the direction of the intrashelf basin (Fig. 4a). These forced regressive wedges are made of good reservoir facies such as grainstone, rudstone and floatstone. They are clinoform bodies which downlap over a few kilometres and may therefore form considerable heterogeneities at the reservoir scale. Subsurface identification. Detailed seismic interpretation and the study of attribute maps along specific stratigraphic horizons are the best way to identify incisions in the subsurface. Comparison of real seismic sections with synthetic seismic sections made from synthetic seismic modelling may help to characterize the incision geometry and fill type (Gre´laud 2005). The identification of incisions from core study is not straightforward. A hint may be provided by the local presence of a sharp silicified and/or dolomitized surface overlain by a lag and then low-energy confined deposits (carbonates or clays), which may be partially or completely dolomitized. These low-energy deposits may grade upward to more open and grainier facies, containing large organisms (rudists, stromatoporoids) or their fragments.

Channels Reservoir properties and creation of heterogeneity. Tidal channels are filled with coarse high-energy facies (Figs 9 & 14). They may then form very good reservoirs or permeability streaks within a reservoir. Alternatively, channels within transgressive systems tracts (early transgression, first flooding of previously exposed platform, Fig. 5) are filled with alternating units of good and poor reservoir properties. The basal surface of channels is not sharp and no silicification along this surface is observed in the field so it should not form a barrier or baffle within a reservoir. Channels are more likely to create high permeability streaks within a reservoir than incisions, if the channel fill is made of high-permeability facies. Prediction. The identification of tidal channel systems is useful to predict the existence of isolated bioclastic sandwaves or connected shoals, which may form very good and extensive reservoirs at the same stratigraphic level as the channels (Fig. 4b). These sandwaves or shoals would be located in the same stratigraphic sequence as the channels, but laterally, in the areas where more subsidence was available, that is, where a thicker stratigraphic sequence is observed. Subsurface identification. Channels are poorly imaged in seismic data since they are filled with very similar facies to the surrounding rocks, which

reduces the impedance contrast at the top and base of the channels. When they are observed on seismic amplitude maps, channels appear as highly sinuous streams. On cores, tidal channels may be characterized by multiple, stacked and highly burrowed basal erosion surfaces and a usually coarse bioclastic fill (rudstone/grainstone). The integration of detailed stratigraphic studies (vertical facies evolution and stacking pattern), facilitated by the comparison with outcrop correlations may also help to identify the intervals corresponding to phases of flooding of the platform and therefore helps to differentiate channels from incisions.

Conclusions A detailed study of the channelized systems which developed on the inner carbonate platform of the Natih Formation of Oman has allowed the differentiation of two types of systems: incisions and channels. Incisions are purely erosive systems with a sharp, single basal surface which develop during phases of exposure of the platform due to a drop of relative sea level, whereas channels are partly erosive but mainly constructive/depositional systems which develop during phases of flooding of the inner platform, due to a rise of relative sea level. Both incisions and channels are filled during transgressive phases, and channels often develop in the topmost part of the incision fill. The main part of the incision fill is made of muddy carbonate facies or clays, corresponding to poor reservoir facies, whereas channels are mainly filled with coarse bioclastic facies, corresponding to good reservoir facies. These elongated channelized sediment bodies create a high-level of lateral heterogeneity within inner carbonate platform deposits at specific stratigraphic levels: below sequence boundaries and at the base and top of transgressive systems tracts, specifically within asymmetrical aggrading sequences of different frequencies (from highfrequency to third-order). These channelized sediment bodies are associated with other types of sediment bodies which develop at the same stratigraphic levels, hence increasing the heterogeneity of these intervals. In fact, high-angle forcedregressive wedges develop in the intrashelf basin on the margins of the platform at the same stratigraphic level as incisions, and isolated sandwaves or larger shoals develop on the inner platform at the same stratigraphic levels as channels, generally in areas where more accommodation is available. The identification of incisions and channels within inner carbonate platform deposits may significantly help stratigraphic correlations, heterogeneity prediction and therefore reservoir production.

CRETACEOUS PLATFORM INCISIONS AND CHANNELS

The results of this study may serve as a generic model for the differentiation of channelized systems within inner carbonate platform successions. They also illustrate the different types and scales of heterogeneities which occur within apparently simple and ‘layer-cake’ depositional systems such as inner carbonate platforms. These heterogeneities are the results of relatively small variations of sea level which lead (1) to the sub-aerial exposure of the platform, (2) to the development and fill of intrashelf basins, and (3) to the onset of higher-energy depositional systems in the overall low-energy inner platform environment. This work presents parts of the results of a joint PhD project between the Carbonate Centre (JVRCCS) in Sultan Qaboos University, Oman and the EGID Institute in the University of Bordeaux, France. We are grateful to the Carbonate Centre (JVRCCS), a SQU-Shell joint venture based in SQU, for the funding of this research. Publication of this article is by the kind permission of Petroleum Development Oman and the Ministry of Oil and Gas of the Sultanate of Oman. We would like to thank the reviewers P. L. De Boer and P. H. Larsen as well as the editor F. S. P. van Buchem for their interesting comments and suggestions which helped to enhance the manuscript. The authors would also like to acknowledge GeoArabia graphic designers A. Egdane and N. ‘Nino’ Buhay IV for their help with the drafting of Figures 4b, 11 & 13.

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Sequence stratigraphy of Cenomanian – Turonian carbonate platform margins (Sarvak Formation) in the High Zagros, SW Iran: an outcrop reference model for the Arabian Plate P. RAZIN1*, F. TAATI1,2,3 & F. S. P. VAN BUCHEM2,4 1

Institut EGID Universite´ Bordeaux 3, 33607 Pessac, France

2

Institut Franc¸ais du Pe´trole, 92506 Rueil-Malmaison, France

3

Present address: NIOC Exploration Directorate, Jomhouri Ave, Tehran, Iran 4

Present address: Maersk Oil Qatar, P.O. Box 22.050 Doha, Qatar *Corresponding author (e-mail: [email protected])

Abstract: A high resolution sequence stratigraphic model has been constructed for the midCretaceous Sarvak Formation (in the High Zagros region of SW Iran) which was deposited close to the eastern margin of the Arabian Plate. The exceptional outcrop quality, displaying the detailed facies patterns in the transition zone from carbonate platform to intra-shelf basin, offers the rare opportunity to distinguish between the relative control of carbonate sediment supply (S) and accommodation (A) on the depositional geometries of third- and fourth-order depositional sequences. Four third-order sequences have been distinguished in the Sarvak Formation, with a duration varying between 1.5 and 3 Ma, and a thickness of 50–150 m. These are in turn composed of fourth- and fifth-order sequences that form the stratigraphic building blocks of this carbonate system. A significant distinction has been made in the third-order sequences between the early transgression (e-TST) when the system was still flat, and corresponds to a ramp setting, and the late transgression (l-TST) when the carbonate platform to intra-shelf basin topography was created. The rate of accommodation creation is identified as the dominant factor controlling the morphology of the depositional profile, and, as such, the driving motor behind the dynamics of this type of carbonate system. The dip angle of the depositional profile has a major influence on: (1) the hydrodynamics of the system; (2) the type of carbonate sediment; and (3) the volume of carbonate sediment produced. A good correlation with the third-order sequences of the Natih Formation in Oman is demonstrated, which supports a dominant control by eustatic sea-level changes and a similar response of the carbonate system to changes in the rate of sea-level rise on the southern part of the Arabian Plate. This outcrop analogue can be considered as a good reference model for the Cenomanian– Turonian carbonate platform margins of the Arabian Plate, but also as a textbook example of the response of carbonate systems to sea-level fluctuations (relative influence of accommodation and sediment supply).

The mid-Cretaceous successions of the Middle East are characterized by very large, shallow-water carbonate platforms within which locally, and at different times, intra-shelf basins developed hundreds of kilometres across and with water depths of 50 – 100 m (Murris 1980). The repeated initiation of these intra-shelf basins shows a typical pattern, that starts with a regional flat platform top, characterized by a muddy facies, which subsequently shows a diversification into a shallow-water platform, generally with a grainy, rudist barrier, and a benthic foraminifera-rich lagoon, and locally organic-rich, muddy intra-shelf basinal deposits (e.g. van Buchem et al. 2002a, b; Davies et al. 2002). In this way, these carbonate systems

represent petroleum systems, where source rocks and reservoir facies interfinger, and a variety of stratigraphic and structural trapping mechanisms can be identified. This paper focuses on one of these intervals, the late Albian/Cenomanian/Turonian, which is known in the Middle East by a different lithostratigraphic nomenclature from country to country (Fig. 1a): shallow-water deposits are known as the Mishrif Formation in the UAE (e.g. Burchette & Britton 1985), the Natih Formation in Oman (Hughes-Clarke 1988; Scott 1990), and the Sarvak Formation in Iran (James & Wynd 1965). The intra-shelf basinal deposits also have different names locally, for example, Shilaif and Khatiyah

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 187–218. DOI: 10.1144/SP329.9 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. (a) Regional lithostratigraphy of the Late Cretaceous in the Middle East (adapted from Burchette 1993 and van Buchem et al. 2002a). (b) Schematic palaeogeographical map of the Arabian Peninsula during the deposition of the Natih Formation (late Albian– early Turonian) (modified after Murris 1980).

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Formations in the UAE and Qatar, and Ahmadi Member in Iran (James & Wynd 1965; Harris et al. 1984; Jordan et al. 1985; Alsharhan & Nairn 1988, 1997; Grantham et al. 1987). The general stratigraphy and palaeogeography of these formations have been studied in considerable detail in the southern part of the Arabian Plate, notably in the United Arab Emirates and Oman (e.g. Murris 1980; Hughes-Clarke 1988; Burchette 1993; Philip et al. 1995). A regional sequence stratigraphic model for this interval was provided by Sharland et al. (2001) and Davies et al. (2002), whereas high resolution sequence stratigraphic studies on the outcrops of the Natih Formation in Oman were published by van Buchem et al. (1996, 2002a). These studies documented the detailed stratigraphic architecture of the shallowwater platform and intra-shelf basinal deposits, and discussed the relative influence of tectonics, eustacy and climate on their formation. Seismic interpretation and seismic stratigraphic modelling of the outcrops confirmed and refined these sequence stratigraphic models (Droste & van Steenwinkel 2004; Schwab et al. 2005; Gre´laud et al. 2006). The only limitation of these studies was that the high-angle prograding geometries, interpreted from seismic, were not exposed in the outcrops of the Oman Mountains. In this paper, we present seismic scale outcrops of the Sarvak Formation in SW Iran that are timeequivalent to the Natih Formation in Oman, and which show several, accessible and fully exposed high-angle platform to intra-shelf basin transitions (Fig. 1b). First, these outcrops are presented and illustrated in detail; second, they are compared to the time-equivalent deposits in Oman; and third, a genetic model is proposed for the formation of the Cenomanian–Turonian platforms on the Arabian Plate.

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boundary of the Arabian Plate. The closure of the NeoTethys during the Neogene was accompanied by a collision of the Arabian Plate and central Iranian Plate, which formed the Zagros Mountain chain (Fig. 2a). SW of the suture zone (the Main Zagros Thrust, MZT) that results from this collision,

Regional geological setting The study area is located in the Zagros Mountain chain in SW Iran, which is the result of a complex geodynamic evolution of the northern margin of the Arabian Plate that started in the mid-Cretaceous and terminated in the Neogene (e.g. Stampfli & Borel 2002). In the Mesozoic, this area belonged to the southern passive margin of the NeoTethys, where a thick succession of shallow-water platforms accumulated. During the Late Cretaceous, a change to an active margin took place with the initiation of a compressional regime that eventually led to the obduction of part of the NeoTethys oceanic crust on top of the continental margin during the Campanian (Ricou 1971; Berberian & King 1981; Alavi 1991). Several phases of deformation succeeded each other in a continuing compressional tectonic regime during the Palaeogene and Neogene, and successive flexural basins formed along the NE

Fig. 2. (a) Location of the study area in the Zagros Mountains. (b) Geological map of the Kuh-e-Landareh anticline in the High Zagros and location of the studied outcrops: Bibi Seydan and Padena (adapted from the Geological Map of Iran at the one million scale): (1) Pliocene to Quaternary; (2) mid to late Miocene (Fars Group); (3) Palaeogene to early Miocene; (4) Campanian to Maastrichtian (Gurpi Fm); (5) Albian to Santonian (Bangestan Group); (6) Jurassic to Early Cretaceous; (7) Pre-Jurassic sequence; (8) Sinandaj-Sirjan structural domain.

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the Zagros is now divided in two large structural zones separated by a main fault, the High Zagros Fault (HZF). In the internal part of the High Zagros (NE), the shortening is expressed by large, southward-dipping thrust planes involving the crust and the Mesozoic and Cenozoic deposits of the Arabian Plate, whereas in the external Zagros Folded Belt (to the SW), the sedimentary rocks are simply detached and deformed in a series of folds with a wavelength in the order of several kilometres (Sherkati & Letouzey 2004). The studied outcrops are located in the Kuh-e-Landareh anticline, 20 km south of the city of Semirom, 150 km south of Esfahan, in the Interior Fars Zone (Fig. 2b). This large anticlinal structure is positioned at the southern border of the High Zagros, in the northern extension of the Kazerun fault, a large transverse fault (directed north –south) cross cuts the different tectonic units of the Zagros chain. The studied stratigraphic interval is mainly the Sarvak Formation (latest Albian to Turonian), which forms the upper part of the Bangestan Group (Fig. 4). The lower boundary of this group corresponds to an unconformity at the top of the Aptian Dariyan Formation. The Bangestan Group consists of the Kazhdumi, Sarvak, Surgah (Laffan equivalent) and Ilam Formations, ranging in age from Albian to Santonian. The facies in these formations vary from shallow-water carbonates to intra-shelf basinal deposits. Only the Laffan Formation is locally terrigenous. The Bangestan Group is overlain by the Gurpi Formation, a deep-water marl that marks the installation of a deep-water environment as a result of the compressional

tectonics controlled flexure of the Arabian Plate margin (Sherkati & Letouzey 2004). The lithostratigraphic nomenclature of the Zagros Mountains has been established by James & Wynd (1965) based on the definition of biofacies that combine an age and a certain environmental notion in a fossil assemblage. Thus the stratigraphic units are by definition diachronous, which explains the difficulty that exists with regional, but also local, stratigraphic correlations, and as a result with the reconstruction of depositional geometries. The regional stratigraphy of the Sarvak Formation has been studied by, for instance, James & Wynd (1965), Setudehnia (1978), Bordenave & Burwood (1990), whereas various aspects of the Sarvak Formation in outcrop and subsurface are reported in a number of confidential OSCO and NIOC reports. Recently seismic and well log based subsurface work on the Sarvak Formation was published by Farzadi (2006) and Taghavi et al. (2007) respectively. Extensive fieldwork in the Anaran area, with subsequent diagenetic and fracturing studies is reported by Sharp et al. (2010).

Material and methods This study focuses on two outcrops of exceptional quality located in the NE flank of the Kuh-eLandareh anticline (Figs 2b & 3). † The Bibi Seydan outcrop, a 1 km long cliff surface that displays a dip transect across a platform to intra-shelf basin transition in the middle part of the Sarvak Formation.

Fig. 3. Location of the studied outcrops and main field sections in relation to the platform margin position during sequence III (Late Cenomanian).

CRETACEOUS PLATFORM MARGIN IN SW IRAN Fig. 4. Overview of the Padena transect: location of the different lithostratigraphic units, depositional sequences, logged sections and major palaeogeographic domains (platform v. intra-shelf basin) as well as detailed photo illustrations are indicated.

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† The Padena outcrop, located 10 km to the east, is a 10 km long cliff surface that shows a continuous outcrop of the full Sarvak Formation, displaying the two margins of an intra-shelf basin. The stratigraphic and sedimentological study of these outcrops is based on the detailed logging of 10 stratigraphic sections (a total of around 2 km), completed with the petrographic analysis of over 1000 thin sections (Taati 2005). Chronostratigraphic information has been provided by the biostratigraphic analysis of over 100 samples for calcareous nannoplankton, benthic microfossils (orbitolinids) and planktonics (Oligosteginid) following standard NIOC procedure (based on the biozonation scheme of James & Wynd 1965). In addition, C and O stable isotope measurements have been made for 80 samples of one of the Padena sections. In principle, the working method is based on the concepts of high resolution sequence stratigraphy established for siliciclastic systems (Posamentier et al. 1988; van Wagoner et al. 1988; Homewood et al. 1992), and further adapted to carbonate systems (e.g. Sarg 1988; Goldhammer et al. 1990; Louks & Sarg 1993; Read 1995; Homewood & Eberli 2000). In this study the following steps are taken: (1)

(2)

(3)

(4)

The first step includes the sedimentological analysis of facies associations (macro- and microfacies) in measured sections, the examination of the depositional geometries at the outcrop scale, and the interpretation of the depositional processes and the succession of depositional environments. The subsequent definition of depositional sequences at different scales is based on the analysis of the vertical evolution of depositional environments, the vertical stacking of stratigraphic units (in particular bedding pattern) and the analysis of depositional geometries. The high resolution stratigraphic correlations established between different sections integrate information about the stacking pattern of high frequency cycles, the physical continuity of stratigraphic surfaces in the field, and the analysis of photo panoramas. These correlations are the basis for the high resolution sequence stratigraphic model. Chronostratigraphic constraints are provided by bioand chemostratigraphic analyses. The facies distribution and depositional geometries in such a sequence stratigraphic model allow reconstitution of the complex evolution of the carbonate system, and in particular the evolution of depositional profiles and the

(5)

depositional processes that accompany the creation of intra-shelf basins. Regional comparison and discussion of the relative influence of the regional and global key factors that control the dynamics of the studied carbonate system.

Principal lithostratigraphic characteristics of the Bangestan Group in the studied zone In the Padena area a complete succession of Albian to Turonian rocks is exposed at outcrop, consisting of the Kazhdumi Formation, the Sarvak Formation, and a chaotic unit at the top that has no formal name (Figs 4 & 5). This succession lies above a stratigraphic disconformity on top of the Dariyan Formation (top Khami Group). The lower part of the Kazhdumi Formation consists of two massive and tabular carbonate units, of 40– 50 m thickness, each of which overlies a more recessive, marly unit with a thickness of about 10 m (Figs 4 & 5). These carbonate units comprise shallow-water platform facies characterized by the abundance of orbitolinids (including Mesorbitolina subconcava), associated with miliolids and calcareous algae. The marly intervals contain a slightly deeper-water fauna with: oligosteginids, the planktonic foraminifera Favusella washitensis, tintinids (such as Colomiella recta) and sponge spicules. This latter faunal association also characterizes the upper part of the Kazhdumi Formation which is represented by a recessive, mostly poorly exposed unit of about 100 m thickness (Figs 4 & 5). Both the benthic and pelagic faunas confirm an Albian age for this formation. This facies evolution is interpreted as a large-scale transgressive sequence, of which the two cycles at the base form the ‘backstepping’ part, followed by a phase of maximum transgression installing an intra-shelf basinal environment in a large part of the Zagros (e.g. Bordenave & Burwood 1990; Bordenave & Huc 1995). The intra-shelf basinal facies change vertically and laterally into a massive carbonate unit which forms the lower member of the Sarvak Formation in this location (Figs 4 & 5). This unit reaches a thickness of 80 m in the central part of the outcrops (sections 1 and 3; Fig. 4); it thins towards the north, has an erosional surface at the top and grades laterally into basinal argillaceous carbonates (section 4; Fig. 4). It is composed principally of grainy bioclastic (rudist) facies, and is organized as large-scale, northwards-prograding clinoforms. This massive unit is overlain by a well-bedded interval of carbonates and argillaceous carbonates 60 m thick, which forms the middle member of the Sarvak Formation (Fig. 5). It is composed of a succession of tabular beds, one to several metres

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Fig. 5. Composite reference section of the Bangestan Group in the Padena area (based on the sections presented in Fig. 9). The Sarvak Formation (late Albian–Turonian) conformably overlies the Kazhdumi Formation and is overlain by a karstified chaotic carbonate unit. It is composed of three sub-units and four third-order depositional sequences (I to IV). The C isotope curve shows a positive shift in the lower part of sequence IV, which is interpreted as the shift caused by the global anoxic event OAE2 at the Cenomanian– Turonian boundary. (1) Daryan carbonate platform (Aptian); (2) mixed argillaceous-carbonate ramp; (3) Orbitolina carbonate ramp; (4) intrashelf basin facies association; (5) outer ramp/base of clinoforms facies association; (6) bioclastic platform margin facies association; (7) back-barrier and rudist-rich open lagoon facies association; (8) inner platform facies association; (9) karstified chaotic unit; (10) regressive facies sequence; and (11) trangressive facies sequence.

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thick, consisting of a muddy facies with benthic foraminifera and thin-shelled rudists typical of an internal platform environment. The upper member of the Sarvak Formation has a thickness of 150 –200 m (Fig. 5), and is composed of several stratigraphic units of varying composition and geometries: (1) massive carbonate beds composed of a grainy facies, rich in rudists and locally organized in large prograding clinoforms; (2), muddy units of the internal platform, with beds several metres thick; and (3) centimetre-bedded and laminated, argillaceous carbonates of the intra-shelf basin. The faunal content of the Sarvak Formation is characterized by an alternation of intervals rich in radiolitid and caprinid rudist debris and intervals rich in benthic foraminifera, locally dominated either by orbitolinids (such as Orbitolina concava, pers. comm. R. Schroeder) or by alveolinids (Praealveolina sp., Ovalveolina) and nezzazatids (Nezzazata sp., Nezzazata simplex, Nezzazata conica). Based on micro- and macro-palaeontological observations a Cenomanian age has been attributed to the main part of the Sarvak Formation, and a late Albian age to the lowest part. The association of the benthic foraminifera Valvulammina sp. –Dicyclina sp. in combination with the presence of Salpingoporella and Taberina bingistani suggest a Turonian age for the upper part of the Sarvak Formation (Biozone 29 of Wynd 1965). This interpretation is supported by chemostratigraphic measurements carried out on section 4 (Fig. 4). The d13C isotope curve shows a positive peak in the transgressive phase of the last sequence (Fig. 5), which has been interpreted as the (weak) expression of the Oceanic Anoxic Event (OAE) 2 that occurred around the Cenomanian –Turonian boundary (Scholle & Arthur 1980; Arthur et al. 1987; Schlanger et al. 1987; Accarie et al. 1996). The upper boundary of the Sarvak Formation in the Padena area corresponds to a very irregular surface overlain by a chaotic unit that locally attains a thickness of about 100 m. This unit consists of monomict blocks of varying size (decimetre to decametre), consisting of limestones that are recrystallized with only very little rudist debris. Locally well-bedded intervals are found, but these are inclined with respect to the overall bedding pattern, and may change laterally into brecciated sediments that form the chaotic facies. No elements have been found to date these deposits, which seem to have a relatively local distribution. The chaotic structure of this unit is interpreted here as the result of karstification and large-scale collapse. However, this interpretation requires further work that should also take into consideration the possibility of uplift at the time of the Kuh-e-Landareh structure, and/or the activation of the Kazerun fault.

In the Bibi Seydan area, about 10 km west of Padena (Fig. 3), only the middle and upper parts of the Sarvak Formation are exposed at outcrop. The middle part here also consists of a succession of flat carbonate beds with an internal platform fauna. The most characteristic element of the upper part is a bioclastic sedimentary body with spectacular several-decametre thick prograding clinoforms dipping with a 308 angle towards the NE, where it passes laterally into intra-shelf basinal deposits (Fig. 6). In the proximal position, the grainy facies interfinger with the muddy shallow-water facies of the platform top. This prograding unit is also overlain by intra-shelf basinal deposits (approximately 20 m), which pass abruptly into a 30 m thick, well-bedded package that contains cryptalgal laminated mudstones at the base. This shallow-water facies is overlain by a heterogeneously brecciated, chaotic type of facies. This unit is possibly time-equivalent to the chaotic unit observed in the Padena area, but seems to be less affected by brecciation and collapse. In the Bibi Seydan area the chaotic unit is overlain by hemipelagic marls with unidentified ammonites. These represent probably the Santonian Gurpi Formation, suggesting that the underlying unit has a Turonian to Coniacian age. The sedimentological and stratigraphic analysis and discussion will now focus on the Sarvak Formation.

Facies analysis The combined outcrop and microfacies study in the Padena and Bibi Seydan areas has led to the proposition of a sedimentary facies scheme with 17 facies, the main characteristics of which are summarized in Table 1. The facies are grouped in facies associations that are interpreted in terms of depositional environments (Wilson 1975; Flugel 1982; Tucker & Wright 1990; Burchette & Wright 1992). Generally speaking, a vertical facies succession is an expression of the lateral migration of different environments that exist along a given depositional profile. In this study, the continuous exposure of palaeo-depositional profiles from platform to intra-shelf basin allows an interpretation of the absolute water depth of the different depositional facies. Five main environments are distinguished; they are described in terms of their facies composition and internal architecture.

I: The inner platform environment (a) Peritidal deposits. Facies I-1 and I-2 are typically found together at the base of decimetre-thick

Fig. 6. Stratigraphic architecture of third-order depositional sequence III (Late Cenomanian) of the Sarvak Formation in the Bibi Seydan outcrop. Four steps are defined in the evolution of this sequence: (1) E.TST, early transgressive systems tract; (2) L.TST, late transgressive systems tract; (3) HST, highstand systems tracts; (4) FRPW, forced regressive prograding wedge. MST, mudstone; WST, wackestone; PST, packstone.

Fig. 7. (a) Overview of the northeastern part of the Padena outcrop (section Padena 5). Note the presence of several erosional surfaces in the lower part of Sequence II, and tabular-bedded carbonate beds of sequence IV that correspond to inner platform deposits. The location of this overview in the transect is shown in Figure 4. (b) Overview of the southwestern part of the Padena outcrop (section Padena 2). Note the complexity of the prograding clinoforms in the margin of the Sarvak platform (Sequences III and IV). The location of this overview in the transect is shown in Figure 4. See also detailed view of the margin in Figure 10.

Table 1. Facies, facies associations and depositional environments in the Sarvak Formation

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platform top cycles. These facies are expressed either as thin layers of mudstone, locally dolomitized and recrystallized, slightly argillaceous and bioturbated, without identifiable fossils (facies I-2), or as millimetre-, cryptalgal- laminated, Lithocodium aggregatum boundstones with miliolids (facies I-1). These facies are interpreted as very shallow-water deposits of the peritidal zone. (b) Restricted inner platform deposits. Facies I-3 consists of locally dolomitized mudstone with chert concretions. Some bioclastic debris is found at the base. This facies passes vertically and laterally into facies I-4, which consists of mudstone to wackestone facies, that is partially dolomitized and characterized by the relative abundance of orbitolinids (with a flat shape, obtaining a diameter of 10– 12 mm) and some bivalve debris. This very shallow-water facies association is observed at the boundary of the lower and middle Sarvak units, where they are deposited at the base of decimetre– metre-scale cycles, which onlap against seaward-dipping erosional surfaces. (c) Open platform top–lagoonal deposits. This facies association comprises three facies, that show an evolution towards more open water conditions. Facies I-5 consists of mudstones and nodular and bioturbated wackestones with benthic foraminifera (miliolids, Nezzazata, Cyclammina, etc.). It is interpreted as subtidal deposits in a relatively confined lagoon. Facies I-6 consists of more grainy packstones and grainstones, rich in miliolids and algal debris, locally associated with the previous facies, and is interpreted as forming in a more open lagoonal environment. Facies I-7 consists of wackestones to packstones, containing miliolids, alveolinids, pelloids and abundant debris of rudists, and is interpreted as an open lagoon environment. These facies, together with the peritidal environment facies described above, tend to be organized in deepening-upward cycles (facies I-1/I-2, I-5, I-6, I-7), which mark a maximum opening of the platform top environment and coincides with a backstep of the platform barrier. Discontinuity surfaces, showing evidence of oxidation and perforation, and laterally continuous at the scale of the outcrops, typically top the platform cycles. These are interpreted as transgressive surfaces overlying minor exposure surfaces, suggesting a very asymmetric architecture of the platform top cycles. The cycles are tabular and have a thickness varying from decimetre to metre scale.

II: Back barrier/outer lagoon environment The back-barrier facies association consists of two facies: facies II-1, characterized by metre-scale packages of bioturbated wackestones to floatstones

with well-preserved rudists, associated with benthic foraminifera (miliolids, Nezzazata, alveolinids, etc.) and facies II-2, which consists of more grainy, tabular layers, consisting of grainstones and rudstones with large rudist fragments. Facies II-2 is commonly associated with dominantly muddy facies, interpreted as the alternation of higher energy deposits in a low-energy background sedimentation (wash-over fans, tidal inlets, etc.). The relatively massive, tabular layers of facies II-1 and II-2 create a sharp facies contrast with the more finely layered, inner platform deposits with which they interfinger. Cycles in the back-barrier environment generally start with a facies succession identical to the one described for the platform top environment above (facies I-1 or I-2, II-3, II-4 and II-5), followed by facies III-1 and III-2. In addition, the back-barrier cycles terminate with a succession of regressive facies that is better preserved than in the platform domain, which makes them more symmetrical.

III: Barrier/platform margin The platform margin and barrier facies association consist of bioclastic-dominated facies, which are deposited in large clinoforms with a depositional angle varying from several degrees to up to 308, and a bedset height varying from several metres up to 50 m (Bibi Seydan; Fig. 6). The internal structure of the platform margin is complex and controlled by the stacking pattern of high frequency cycles. In a basinward direction, an interfingering with the intra-shelf basinal facies is observed, individualizing metre-scale, shallowing-upward cycles, that consist of the following facies succession: at the base, facies III-3 occurs (very fine-grained bioclastic packstones and grainstones, that are generally nodular and bioturbated and rich in rudist and bivalve debris). This facies forms relatively thin and subhorizontal layers, and is interpreted as the bottom sets of the clinoforms. It is overlain by more massive, coarse-grained grainstones of facies III-2, which have little bioturbation, and are composed of rudist debris and rare benthic foraminifera. These deposits are locally parallel bedded to the overall clinoform shape. The upper part of the cycles, facies III-1, is generally very coarse-grained, consisting of bioclastic grainstones and rudstones with well-preserved rudists, associated with a mix of benthic fauna, including debris of algae, bivalves and various foraminifera (Nezzazata, alveolinids, miliolids). These deposits are organized in decimetre- to metre-scale, sub-horizontal, slightly inclined, seaward-dipping beds. The top surface of each cycle is generally bioturbated. The organization of the cycles shows the progradation of the platform top towards the intra-shelf

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basin. The dimensions of these bioclastic bodies (thickness and lateral extent) depend on the stratigraphic context, and both are controlled by the ratio of accommodation to sediment supply. Along the dip direction, they vary at the hundred metres to kilometre scale, and in the strike direction dimensions are likely to be at the multi-kilometre scale.

IV: Intra-shelf basin environment The intra-shelf basin facies association consists of a monotonous alternation of decimetre-scale, tabular beds that consist of three types of facies. Facies IV-1, consisting of nodular and bioturbated wackestones and packstones with fine-grained bioclasts including debris of rudists and diverse bivalves, is typically found at the toe of the slope, the most distal part of the bottom sets of the clinoforms. This facies is locally developed into thicker packages, intercalated in the more basinal facies, and is interpreted as a lowstand deposit. In one of these lowstand packages a unique occurrence of very fine-grained, siliciclastic fraction has been observed. Facies IV-2 is characterized by a vaguely expressed alternation of decimetre-scale beds, consisting of nodular mudstone with planktonic foraminifera (globigerinids), oligosteginids (calpionellids) and sponge spicules. It occurs in the more distal domain, and alternates with facies IV-3. Facies IV-3 consists of millimetre-scale, laminated mudstones with rare fossil fragments and frequent, flattened, silica nodules on the bedding planes. Facies IV-1 and IV-2 are dark coloured, rich in silica nodules, and lack bioturbation. The general depositional environment of this facies association is interpreted as a low-energy, intra-shelf basin setting, with a water depth in the order of 50 –70 m, as indicated by the observed geometries in the platform to basin transition.

Sequence stratigraphic model: geometries and facies distribution of the Sarvak depositional sequences The stratigraphic architecture of the Sarvak Formation has been reconstructed combining the bedding pattern and depositional geometries with a detailed macro- and microfacies analysis of measured outcrop sections. Correlations were established by walking along special surfaces (hardgrounds, firmgrounds and erosional surfaces) and from the interpretation of photo panoramas. The high resolution, time-based, stratigraphic framework thus obtained, in combination with the analysis of the stacking pattern and sedimentary facies, represents a unique database to illustrate the dynamics of this carbonate sedimentary system.

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Four main depositional sequences are distinguished in a time interval spanning the late Albian to the early Turonian (Fig. 5). With an average duration in the order of 2–3 Ma, these sequences fall in the range of third-order depositional sequences sensu Vail et al. (1991). These are composed of several fourth-order sequences, which themselves are composed of metre-scale fifth-order cycles. All four third-order sequences are well exposed in the Padena area, but only the two last sequences are accessible in the Bibi Seydan area. In order to present a model for the evolution of the depositional sequences in the most coherent way, we will first present the detailed analysis of one third-order depositional sequence that is spectacularly exposed along the cliffs of the river valley in the Bibi Seydan area. This is followed by a presentation of the complete system composed of four third-order sequences in the Padena area.

Bibi Seydan outcrop In the Bibi Seydan outcrop, Sequence III is exposed perpendicular to the coastline, displaying the full geometrical complexity of this platform to intrashelf basin transition over a distance of 1 km in the two opposing valley walls (Figs 3 & 6). The quality and accessibility of the outcrop allows studying the sedimentary facies in the framework of the depositional geometries. This offers the unique opportunity to propose an unambiguous palaeobathymetric interpretation of the sedimentary facies, and thus serves as an excellent introduction to the stratigraphic organization and facies model of the Sarvak Formation. Three geometrically different stratigraphic units can be distinguished: (1) a basal unit, consisting of horizontally-bedded, intra-platform facies; (2) a middle unit, that reaches a maximum thickness of 90 m, and shows a lateral facies and bedding change from horizontally-bedded intraplatform facies, to large scale, bioclastic clinoforms that interfinger in the distal position with intra-shelf basinal facies, and (3) an upper unit, consisting of deeper-water intra-shelf basinal facies (Figs 6 & 8). By comparison with the Padena outcrop, the first two units are of Cenomanian age, and the Cenomanian –Turonian boundary is located in the upper unit. The basal sequence boundary of Sequence III is positioned 30 m below the top of the horizontallybedded intra-platform facies of the lower unit, in an interval where the platform cycles have a minimum thickness and are mainly composed of peritidal facies (facies I-1 and I-2; Fig. 6). This facies is interpreted as marking a period of minimum accumulation rate on top of the platform, and the sequence boundary corresponds to one of

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Fig. 8. Sequence stratigraphic model for a third-order sequence in the Sarvak Formation (based on the Bibi Seydan outcrop). Note the progressive increase in the inclination of the depositional profile (clinoforms) during the aggradational phase of the carbonate platform and the time-equivalent development of high energy bioclastic deposits replacing the initial muddy ramp facies. (1) E-TST, early transgressive systems tract; (2) L-TST, late transgressive systems tract; (3) HST, highstand systems tracts; (4) FRPW, forced regressive prograding wedge.

the many exposure surfaces observed in this interval (see below). The upper boundary of Sequence III is represented by an erosional surface which is abruptly overlain by intra-shelf basinal deposits. This surface is slightly dipping towards NE (basinward direction) and truncates the underlying clinoforms, creating a toplap contact, in the northern part of the outcrop (Fig. 6). Locally, it forms a succession of small downward-stepping terraces. To the south, in the proximal domain, the sequence boundary changes to a concordant horizontal surface without any evidence for erosion. The toplap surface is interpreted as the result of a marked progradation of the platform (forced regression). It was probably slightly modified by the ravinement during the following transgression, but the amount of erosion linked to this event was probably very limited indeed, as can be concluded from the virtual absence of reworked material along this surface. Sequence III, estimated to be of late Cenomanian age, is positioned in an overall transgressive trend, between Sequence II (middle Cenomanian?) consisting entirely of shallow-water platform facies, and Sequence IV (late Cenomanian –early Turonian) consisting, here, of intra-shelf basinal facies. Based on the depositional geometries and the sedimentary facies four systems tracts have been distinguished, which correspond to the four principal stages in evolution of the deposition of this sequence (Fig. 8). The first stage is represented by a tabular and semi-horizontal basal unit consisting of highfrequency platform top cycles bounded by subaerial exposure surfaces. In detail, the lower cycles have a metre-scale thickness and are dominated by a very shallow-water mudstone facies with Lithocodium and Miliolids (facies association I-1, I-2, I-3) whereas the upper cycles are thicker (several metres) and consist of wackestone to

floatstone facies with alveolinids and rudists typical for an open lagoon (facies I-6, I-7 and II-1). This evolution is interpreted as the uniform aggradation of an internal platform over which the carbonate production (S) was equal to the rate of accommodation (A). The thickening-upward trend observed in the cycles, in combination with the facies evolution suggest a gradual increase in the rate of accommodation, typical for the ‘early transgressive systems tract’ sensu Vail et al. (1991). The second stage is characterized by the progressive development of NE-dipping clinoforms in a generally aggrading system. These have a sigmoidal shape, and appear above the tabular basal unit, initially with a very low angle (,18), that gradually increases to a very steep angle of 308 (Figs 6 & 8). The evolution of the depositional geometries, and the concomitant depositional profile of the carbonate system, is accompanied by a lateral differentiation of the facies along the clinoforms, as well as a gradual change of this system through time. The first low angle clinoforms consist of muddy facies of wackestone and floatstone with benthic foraminifera and rudists in the proximal domain, and of intensely bioturbated wackestones and packstones with benthic foraminifera and bivalve debris in the distal domain. The intensity of the bioturbation and the relative concentration of bioclastic material reflect the condensation that occurs in the distal, thinning part of the clinoforms. During the gradual growth and geometrical change of the clinoforms a clear facies differentiation occurred. In the proximal domain a vertical aggradation is observed of the intra-platform cycles (facies association I) with the regular intercalation of back-barrier facies rich in well-preserved rudists (facies association II). At this stage, the highfrequency cycles are several metres thick and thicken-upward. In the margin domain, high-energy

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bioclastic rudist facies are deposited in clinoforms with an increasing dip angle (20–308). The grainstones and rudstones of the upper part of the clinoforms (facies III-1 and III-2) pass laterally on the slope into finer-grained grainstone and packstone, and subsequently into mudstone and wackestone facies with planktonic foraminifera characteristic of the intra-shelf basin environment (facies association IV). The distribution and type of facies change significantly with the variation of the topography. The grainy facies belts, which represent the bioclastic barrier, tend to widen by spreading out laterally towards the platform into the lagoon, as well as into the intra-shelf basin. At the same time, at the base of the clinoforms, the bioturbated wackestones evolve into black millimetrelaminated mudstones, without bioturbation and with a fetid smell (facies IV-3). The depositional profile along the clinoforms suggests a water depth for the confined intra-shelf basin of approximately 40– 70 m. One of the most important observations of this study is that the geometrical turn-around point (the change from aggradation to progradation) occurred while the bedding pattern (thickening upward) and facies evolution in the intra-platform indicate that the accommodation was still increasing. From this it follows that the geometrical expression is the result of an increase in the production of the carbonates, which overtook the increase in accommodation before it had reached its maximum. In this situation, the concept of the maximum flooding surface becomes a delicate matter (see discussion below). The third stage is characterized by a progradational trend with clinoforms evolving from sigmoidal to more oblique and with a toplap termination surface. The platform top cycles gradually become thinner, and finally disappear. This geometrical evolution suggests a gradual decrease of accommodation to zero, and the system only prograded with no further sediment accumulation on the platform top. The carbonate system was then reduced to a platform margin wedge, consisting essentially of coarse-grained bioclastic material with rudists (facies V-2) that accumulated on large prograding foresets with a topographic relief of 70 –80 m and an inclination of 308. At the base of the clinoform, layers of packstone and fine-grained grainstone (facies III-3) alternate with muddy intra-shelf facies. No evidence for gravity flow deposits has been found. This stage corresponds to the highstand systems tract of Sequence III (Fig. 8). During the fourth stage, the progradation of the bioclastic wedge continued, but the toplap surface at the top of the clinoforms was inclined towards the basin. The facies distribution is identical to that described for the highstand phase, but the geometry of the clinoforms is significantly different (Fig. 8)

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with a diminution of the inclination and a thickening of the deposits at the base of the clinoforms. These combined observations suggest a fall of relative sealevel below the platform top (negative accommodation), and a shift of the depocentre towards the intra-shelf basin. The deposits of this stage are interpreted as a forced, regressive, prograding wedge (sensu Hunt & Tucker 1992, 1993). The relative sealevel fall is estimated in the order of 15–20 m, depending on the amount of erosion related to the ravinement surface (estimated to be little). The bioclastic facies of Sequence III is directly overlain by intra-shelf basinal facies, indicating a rapid backward step of the depositional system in the next sequence in this area. In the absence of clear unconformity and downward shift of facies in the prograding part, the upper boundary of Sequence III is placed at the toplap surface, that is above the forced regressive deposits (van Wagoner et al. 1990; Hunt & Tucker 1992), and not at the base of the forced regressive deposits, such as proposed by Kolla et al. (1995). No lowstand deposits have been found here contrary to what has been observed in the Padena area. The clear expression of the architecture of Sequence III in the Bibi Seydan location unambiguously demonstrates that the geometry of the depositional profile and the dynamics of the carbonate system, and thus the spatial distribution of the facies in a sequence, are principally controlled by the interaction of the evolution of the rate of accommodation (A) and the carbonate production (S). This example is therefore a perfect illustration of the theoretical model proposed by Schlager (1991, 1993, 2005) and Homewood (1996). A reference model for the evolution of a carbonate system during the deposition of (large scale) depositional sequences can be summarized in four steps that make up an accommodation cycle (Fig. 8): (1) Early transgression After a period of sub-aerial exposure, the gradual rise of relative sea-level creates the conditions for the deposition and preservation of a package of very shallow-water sediments deposited along a sub-horizontal depositional profile. This is a low energy system, characterized by muddy sediment in an intra-platform setting. At this stage the carbonate production follows after the creation of accommodation space (A ¼ S), and an aggrading stacking pattern is formed. The slow overall increase in accommodation rate is expressed by a gradual thickening-upward trend in the platform cycles. The deposits of this stage represent the early transgressive systems tract. (2) Late transgression The acceleration in the rate of relative sea-level rise causes, at the scale of the platform, a deficit in the carbonate

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production with respect to the creation of accommodation (A . S). When this deficit is homogeneous at the scale of the platform, a general deepening is observed and eventually a drowning of the platform (Schlager 1989). In this case study, the production deficit is expressed by a non-homogeneous sedimentation rate at the scale of the platform, which leads to a differentiation of shallow-water areas, where the carbonate production keeps up with the relative rise in sea-level (A ¼ S), and deeper-water areas, also called intra-shelf basins, where the carbonate production does not keep up with the rising sea-level and where local to regional bathymetric lows are formed (A . S) through the differential sedimentation rate. This is the fundamental process that underlies the origin of much of the submarine palaeotopographical relief in carbonate systems, but also of the creation of clinoforms. It entails a profound modification of the dynamics of the platform and of the depositional system. The increase in water depth favours the influence of high-energy sedimentary processes, such as wave action and tidal currents, which create bioclastic barriers at the top of the clinoforms. The high energy environment has the tendency to distribute sediment laterally during the growth of the clinoforms. In other words, a bioclastic barrier-rimmed platform and intra-shelf basin replace the muddy ramp system of the lower part of the sequence. This change in the sedimentary system is accompanied by an increase in the carbonate production rate, to such an extent that it causes a progradation of the system while the accommodation is still increasing (A , S). In terms of definitions, the change from an aggrading trend to a prograding trend is interpreted as the turnaround point (A , S) in an ‘A/S’ cycle, that is, the ‘maximum flooding surface’ that separates the TST and HST in sequence stratigraphic terms. In an ‘A’ cycle, this maximum flooding surface (MFS) lies much higher in the sequence, and is much more difficult to pick in the platform margin, contrary to the situation in the platform interior, where the accommodation control is the ruling parameter. The difficulties that can arise when correlating these domains in less ideal situations are obvious (see also discussion below and Fig. 16). (3) Early highstand The slowing down and subsequent stop of the relative rise in sea-level is expressed in the sediment as a pure progradation without an aggradational component (A , S), and the evolution from large, sigmoidal clinoforms to smaller, oblique clinoforms with a steep angle. The bioclastic deposits and

their equivalent deposits on the platform represent the highstand systems tract. (4) Late highstand Finally, the fall in relative sealevel causes the deposition of a forced regressive wedge characterized by less inclined clinoforms and an upper toplap surface that dips gently seawards. On the platform top this phase corresponds to an emersion surface. The upper sequence boundary may be affected by the transgressive ravinement surface of the next sequence.

Padena outcrop The vertical and lateral continuity of the outcrops in the Padena area was used to construct a sequence stratigraphic model for the four sequences that compose the Sarvak Formation (Fig. 9). Sequence I (late Albian –early Cenomanian). Only the upper part of Sequence I is exposed in the Padena area. It corresponds to the lower Member of the Sarvak Formation, and consists of a massive carbonate body of bioclastic rudist debris with a maximum thickness of 80 m (field section 3; Figs 4 & 9). The base of this unit consists of an alternation of sub-horizontal, metre thick beds that constitute high frequency shallowing-up cycles. The facies succession in these cycles starts at the base with nodular and chert-rich mudstones and wackestones, with fine grained bioclasts and micro planktonics (facies V-1), overlain by fine-grained bioclastic packstones (facies III-3). The next facies is a well-developed bioclastic grainstone package that shows a coarsening-upward trend and is characterized by large foresets dipping at an angle of 20 – 308 towards the ENE (facies III-2), interpreted as a platform-margin environment. The top of the clinoforms is a minor toplap surface that indicates a downward shift of the facies belts since it is overlain by tabular beds of rudist grainstones and floatstones, interpreted as back-barrier and platform top deposits (facies III-1 and II-2; Fig. 9). These deposits pass laterally towards the ENE into grainstones facies resembling those found at the base of this sequence. Towards the SW (field section 1; Figs 4 & 9) a lateral change to a muddy facies with rudists and benthic foraminifera of the intra-platform is observed (facies I-5, I-6, I-7, II-1). The sequence boundary at the top of this succession is represented by a major discontinuity surface, which also marks a very abrupt facies change. The evolution of the facies and geometries of the beds are interpreted as the result of a prograding carbonate system, with a sigmoidal depositional profile and characterized by a rudistic rich bioclastic platform margin. The observed bathymetric profile from the platform top to the base of the clinoforms was in the order of 50 m.

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Fig. 9. Sequence stratigraphic model for the Sarvak Formation in the Padena outcrop based on logged sections and lateral following out of beds. Four third-order sequences are distinguished, which are subdivided in medium scale (fourth-order) sequences. Note the creation of an intra-shelf basin in Sequence III on top of the platform of Sequence II. This topography was mainly created through differential sedimentation, and not by tectonic deformation (see text). Figure 4 shows a panoramic photo of this outcrop, and detailed illustrations of the outcrops are presented in Figures 7a, b and 10.

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Sequence II (Cenomanian). The lower boundary of Sequence II is represented by a major discontinuity surface that corresponds to a NE (seaward) inclined erosional surface at the platform margin (between sections 3 and 4; Fig. 9), and passes towards the SW (landward) into a concordant surface marking a sharp facies change and a change in the stacking pattern within the intra-platform environment (between sections 1 and 3; Fig. 9). Generally speaking, this depositional sequence consists of a vertical stacking of metre-scale beds composed of internal platform facies (facies association I). This layer cake architecture has been observed at the outcrop scale, distinguishing this sequence from the more complex architectures of Sequences I and III, which developed in intra-shelf basinal depressions (Fig. 9). An overall NE thickness increase of Sequence II is observed, from 50 m thick in sections 1 and 3, to more than 100 m in sections 4, 5 and 6 (Fig. 9). This thickening occurs particularly in the lower part of the sequence. The analysis of the facies and bedding pattern shows that Sequence II consists of about ten higher frequency cycles (probably fourth-order), with a thickness in the order of 5–10 m. These cycles are asymmetrical and are characterized by a thickening-upward trend of smaller scale cycles, that are composed of very restricted lagoonal facies at the base (facies I-1 and I-2), followed by a more open water facies (facies I-5 to I-7). In the upper part of the intra-platform cycle a thin succession of thinning-upward beds can be preserved, consisting of restricted, bioturbated lagoonal facies (facies I-5), interpreted as the platform expression of the regressive phase (loss of accommodation). This regressive phase is very poorly developed or not present at all, since in the open lagoonal environment the facies (facies I-6 and I-7) are directly overlain by a hardground or firmground, making asymmetrical cycles. Hardground surfaces that cap these cycles, are continuous at outcrop scale and are interpreted as exposure surfaces, providing an excellent high resolution time framework. In detail, the lateral facies distribution within the cycles is relatively complex, and does not show a particular polarity. A unique feature of Sequence II is given by the three slightly inclined erosional surfaces found in the north-eastern part of the transect, in onlap position against the topography of the previous sequence (between sections 3 and 5; Figs 7a & 9). These surfaces have similar geometries, and are covered by similar types of sediment. The first one corresponds to the lower boundary of Sequence II in this area, and the other two are located in the lower part of the same sequence and can be correlated towards the SW with hardground surfaces that correspond to the top of high frequency cycles

(second and fifth high frequency cycles). These surfaces are covered in onlap position by an unusual facies consisting of mudstones with cherts (facies I-3) and mudstones with large orbitolinids (facies I-4), interpreted as indicators of shallow and confined environments. These deposits evolve rapidly into massive, metre-scale beds composed of grainstone and bioturbated floatstone with rudists, chondrodonts and benthic foraminifera (facies II-2), that cover the platform towards the SW (Figs 7a & 9). At the top, these cycles are bounded by sharp, locally erosive discontinuity surfaces. This type of facies association is interpreted as corresponding to a low-energy, intra-platform environment followed by a higher-energy environment and as deposited along a gently seaward-dipping depositional profile in an overall onlap and backstep stratigraphic position formed during early transgression. The erosional surfaces are presumed to result from drops in relative sea-level in the order of 20 –40 m, deduced from the stratal relationships. These surfaces mark the base of lowstand packages of high frequency cycles deposited in a shelf margin wedge. The fact that the cycles are not eroded in the distal position, in combination with the aggrading pattern of the deposits in which these incise, suggests a deposition in an overall increase of accommodation, rather than in a forced regressive context. The slight progradational nature of the position of the erosional surfaces is explained by the progradation of the shelf margin wedge. Such evidence of sea-level drops at a similar stratigraphic level has been inferred from incision features within the Natih carbonate platform in Oman (van Buchem et al. 2002a; Gre´laud et al. 2006, 2010). Above the top erosional surface, Sequence II comprises stacked, tabular, high frequency cycles. All of these cycles thicken towards the NE (Fig. 9). The general facies distribution is independent from this thickening. The most open water, lagoonal facies is observed in the middle part of the transect, and is thus located above the platform margin of the underlying Sequence I, and below the axis of the intra-shelf basin that developed in Sequence III. The lower part of this package (Fig. 9), is relatively thick and contains a maximum of diversity of benthic foraminifera, which has been used as the criterion to position the MFS of Sequence II here. Twenty-five to forty metres above this MFS, relatively thin, high frequency cycles, mostly consisting of peritidal facies (facies I-1 and I-2) separated by surfaces of subaerial exposure, indicate a reduction in accommodation on the platform top, and are interpreted as the top boundary of Sequence II. Sequence II, as exposed in these outcrops, is thus formed by an overall vertical aggradation of intra-platform cycles in response to a general slow

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increase of accommodation. In this context, the carbonate production compensated the variations in accommodation, by following closely the increasing sea-level, and keeping a virtually horizontal depositional profile. High frequency periods of decrease in accommodation caused successive phases of sub-aerial exposure. This is particularly evident in the lower part of the sequence, in the platform margin wedge of the lowstand, where three, slightly seaward-dipping erosional surfaces have been observed that indicate relative sea-level drops of the order of 20–40 m. The facies associations in these cycles consists of low-energy orbitolinid facies (facies I-3) and high-energy rudist facies (facies II-2), whereas the confined lagoonal facies (facies I-3 to I-5) is present only at the platform top. The general thickening of Sequence II towards the NE is interpreted as the result of a regional differential subsidence, compatible with the platform margin location of Sequence I and the position of the shelf-margin wedge. This differential subsidence is, however, continuously compensated by the carbonate production, and has not influenced the facies distribution on the platform. The position of the most open water facies in the middle part of the transect reflects a morphological influence that was independent from the subsidence, but could be related to the position of the platform margin of the underlying sequence. Sequence III (mid? to late Cenomanian). Sequence III represents a facies succession and depositional geometries that are very similar and time equivalent, to those described in detail for the Bibi Seydan outcrop (see above; Figs 6 & 8). Here we will only summarize the main characteristics. At the base, the sequence consists of a very flat bedded unit with intra-platform facies (facies I-5 to I-7), which is interpreted as an early transgressive systems tract, constructed by vertical aggradation in an overall slow rise in sea-level. In the southwestern part of the transect it is overlain abruptly by a relatively condensed interval of intra-shelf basinal facies (facies association IV), which to the NE passes into shallow-water platform deposits organized in clinoforms that show a gradual increase depositional angle (Fig. 9). This interval corresponds to the phase of intra-shelf basin creation, and is controlled by differential aggradation of the platform at the time of maximum rate of accommodation creation. The upper part of the sequence consists of a prograding bioclastic platform margin unit (facies association III) and rudist-rich back barrier deposits (facies association II) organized in sigmoidal or oblique clinoforms with a depositional angle of up to 208 (Padena 5 section; Figs 8 & 9). The prograding, bioclastic marginal deposits form massive,

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grainy carbonate bodies of more than 70 m thickness, with a width of about 2–3 km in dip direction. Their lateral continuity, along strike, is likely to be at the scale of tens of kilometres (Fig. 3). This aggrading and prograding system results from a sediment production that was greater than the rate of accommodation. The upper sequence boundary delineates an abrupt downward shift of facies, and corresponds to a toplap surface at the top of the clinoforms in the platform margin (sections Padena 2 and Padena 4; Figs 7b & 9). In the platform margin setting of the Padena 2 section, a carbonate lowstand package consisting of shallowwater facies is deposited in onlap position against this sequence boundary, approximately 15 m below the shelf edge (Fig. 7b). This sequence boundary is thus interpreted as a major surface, marking a period of sub-aerial exposure as a result of a relative sea-level drop in the order of 10 m. A unique aspect of the Padena outcrop is that two opposite margins of the intra-shelf basin are exposed (Figs 4 & 7), which, in combination with several spot observations made in the region, allows construction of a palaeogeographical map (Fig. 3). This map illustrates that the platform margin had a sinuous shape, and a complex organization of the clinoform belt is inferred. It also illustrates that the intra-shelf basin was not completely filled during sequence III. The characteristics of Sequence III observed in the Padena area correspond well with the model proposed in the Bibi Seydan area for the evolution of the depositional system and the processes instrumental in the creation and infill of an intra-shelf basin (Fig. 8). These outcrops confirm the presence of a fall in relative sea-level at the top of the sequence, that stopped the carbonate production and with that the progradation of the margin. A key question in the understanding of the creation of intra-shelf basins is what controlled the localization of these basins and the zones where shallow-water carbonate production kept up with sea-level. In the Padena area, it appears that the intrashelf basin is not located in the area with the strongest subsidence. On the contrary, the area of maximum subsidence located in the north-eastern part of the transect corresponds to a zone of maximum aggradation of the shallow-carbonate platform deposits of sequence II and III (Fig. 9). In this case, the localized increase of subsidence favoured carbonate production, constraining the area where the shallow-water carbonate platform kept up, maintaining a slight high during deposition of sequences II and III as is shown by the facies (Fig. 9). This minimal topographic relief may have been sufficient to control the carbonate production during the increase in rate of relative sea-level rise, and thus the area of aggradation bordering the intra-shelf

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basin. This hypothesis requires further investigation in the three-dimensional organization of the system. Sequence IV (late Cenomanian– early Turonian). Sequence IV starts with a lowstand wedge in onlap position against the lower sequence boundary (Figs 7 & 10; section 2), that consists of parallel-bedded layers of wackestones and bioclastic packstones and a terrigenous fine fraction of quartz. This siliciclastic influence is unique in the Sarvak, and stresses the importance of this sequence boundary. These facies pass laterally into intra-shelf basin facies. On the platform, the basal sequence boundary corresponds to a sharp and laterally continuous horizontal surface that coincides with the transgressive surface. Across this surface a downward shift of the facies belts occurred since intra-platform deposits are directly deposited on top of the prograding bioclastic facies of the underlying sequence (Figs 7 & 11). However, this downward shift was not accompanied by a migration of the bioclastic barrier, but rather by its disappearance. During the next sea-level rise it was again gradually reconstructed. This observation illustrates a fundamental difference with the siliciclastic systems, and is explained by the decline and virtual cessation of carbonate production due to sub-aerial exposure. On the platform, this sequence consists of a succession of four (fourth-order) cycles, varying in thickness from 15–35 m (Fig. 11). In each one of these cycles a lateral facies change can be observed, over a distance of 2 km, from the internal platform association (facies associations I and II) to bioclastic platform margin facies, and, in the most distal position, to intra-shelf basinal facies (Fig. 11). The

platform margin remains roughly at the same geographical position as in the previous sequence (Sequence III), and consists of bioclastic sedimentary bodies with a thickness of approximately 20 m, and a dip-direction width in the order of 1 km. In detail, the stacking pattern of these bioclastic bodies shows a slight retrogradation that occurs around the Cenomanian –Turonian boundary (cycles IV-2 to IV-4). This slight landward migration of the system, in combination with the preservation of the intra-platform deposits, occurred during relative sea-level rise. Contrary to the Bibi Seydan area, where this transgression led to the drowning of the carbonate platform and an abrupt back step, the increase in accommodation is here compensated by a strong increase in the carbonate production rate causing an aggradational trend, and only a very small back step of the platform. This difference is a fine illustration of the heterogeneous character of the carbonate production rate in transgressive phases, and the different effect it can have on the resulting depositional geometries along the same coastline. The third-order transgressive systems tract is truncated by a stratigraphic discontinuity marking the base of the chaotic unit (Fig. 7a). The complex and chaotic nature of this top unit is mostly likely caused by large scale karstification and collapse. The preceding sea-level drop probably occurred after the early Turonian and before the Santonian transgression, leading to exposure and terrestrial conditions (Laffan Formation), as has been reported from other places in the Zagros by Wynd (1965). Medium scale sequences in large scale Sequence IV. The exceptional quality of the outcrops between

Fig. 10. Detailed view of the southwestern part of the Padena outcrop, illustrating the geometries of the prograding bioclastic deposits in Sequences III and IV. Note the position of the lowstand wedge (medium scale Sequence IV.1) onlapping the clinoforms top of the underlying medium scale Sequence III.2. The medium scale sequences IV.2, IV.3 and IV.4 show a retrograding, aggrading to prograding stacking pattern (see location of this detailed view in Fig. 7b).

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Fig. 11. Detailed sequence stratigraphic model consisting of third- and fourth-order depositional sequences illustrating the architecture of the platform to intra-shelf basin transition. This transect has been established by physical correlation between sections Padena 1 and 2 (Figs 4, 7b & 10). The reconstitution demonstrates (a) the gradual increase of the depositional profile inclination within fourth-order depositional sequences, (2) the alternation of low energy muddy depositional systems during early transgressive stages (E-TST), and high energy, grainy depositional systems characterized by bioclastic platform margin deposits during late transgression (L-TST) and highstand (HST), and (3) the volumetric partitioning between transgressive systems tracts best preserved on the platform top, and the regressive deposits best preserved at the platform margin.

sections 1 and 2 (Figs 7b & 10) means that the beds in Sequence IV can be followed, creating a highresolution sequence stratigraphic framework within which geometries and lateral facies changes can be documented in detail from the internal platform to the intra-shelf basin. All four highfrequency cycles show a similar facies pattern (Fig. 11), that strongly resembles the one observed at the third-order scale, described above for the Bibi Seydan outcrop (Fig. 8). At the base, the cycles start with very shallow-water lagoonal facies deposits (facies I-1, I-2, and I-5) which change laterally to slightly more open-water platform deposits with benthic foraminifera (facies I-5, I-6). The layers are thin, tabular, and wedge out laterally, towards the margin. At this stage, no high-energy grainy facies is formed on the platform. The dominantly muddy carbonate sediment is mostly of microbial and algal origin, with some admixed shell material of benthic foraminifera. This facies association represents the early transgressive systems tract. It is followed by a differentiation of the sedimentation rate between the internal platform and the margin, initializing the growth of the basinwarddipping clinoforms (Fig. 11). This geometrical evolution is accompanied by a change in the nature and

the distribution of the sedimentary facies. In this stage the grainy barriers (facies association III) and the associated back-barrier rudist deposits (facies association II) develop gradually. The facies distribution and evolution show a clear retrogradational pattern, as well as a widening of the barrier which reaches a maximum width of 500– 1200 m just below the maximum flooding surface. The time-equivalent internal platform deposits consist of back-barrier facies with abundant rudists. This late transgressive systems tract is well developed on the platform top, but thins towards the margin, and corresponds to a condensed surface in the intra-shelf basin (Figs 7b, 10 & 11). The transgressive systems tract is characterized by a gradual change in the carbonate system from (1) a slight basinward dip of the depositional profile, created by a high microbial carbonate productivity (algal, bacterial) in the internal part of the platform; (2) the gradual development of more grainy sediment production in the high energy, external part of the depositional profile; and (3) the expansion of the rudist-rich facies, that eventually cover a very large area behind the bioclastic barrier that is being created. The maximum flooding period is characterized on the platform top, in the internal domain, by

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thick layers of back-barrier facies (grainstones, floatstone with rudists), that pass laterally into bioclastic facies, organized in aggrading clinoforms, that wedge out in a condensed level in the intra-shelf basin (Fig. 11). The regressive part of the sequences is thin in the internal platform domain, and consists of thin beds with a thinning-upward trend, that have a similar rudist-dominated facies as the underlying transgressive sequences. These are overlain by a sub-aerial exposure horizon, at the cycle level. These deposits thicken towards the basin centre, and change laterally to grainy, bioclastic deposits organized in sigmoidal clinoforms along the margin of the platform. This bioclastic barrier facies pass distally into intra-shelf basin sediments deposited in approximately 30 –40 m of water depth. Subsequently, the platform prograded in oblique clinoforms that have a toplap contact with the overlaying sequence boundary (Fig. 11). The top surface of the clinoforms steps down towards the basin, implying a forced regressive trend in the final part of the sequence, in response to a fall in relative sealevel. This forced regressive trend is characterized by a reduction in the size and angle of dip of the clinoforms, caused by a sharp drop in carbonate production during exposure of the platform. The stratigraphic organization of these mediumscale cycles shows a clear volumetric facies partitioning between the transgressive phase, represented by a strong aggradation of the internal platform and starvation of the intra-shelf basin, and the regressive phase, consisting of a bioclastic progradation platform wedge, initially with sigmoidal clinoforms, and later with oblique clinoforms. The architecture of the various high-frequency cycles that make up the transgressive systems tract of Sequence IV, is identical. Apart from the backstepping trend, the principal evolution observed consists of the thickening upward-trend in the intraplatform deposits, and the forced progradational wedges along the platform margin, and evidence for onlapping lowstand packages in cycles 4.1 and 4.4 (Fig. 11). This evolution is interpreted as an increase in the amplitude of the high frequency relative sea-level variations around the Cenomanian– Turonian boundary. The evolution of the carbonate system as deduced from the depositional geometries and the facies distribution within the high frequency cycles appears to be very similar to the one proposed to explain the architecture of the third-order sequence at Bibi Seydan and Padena (compare Figs 6 & 10). This observation supports the hypothesis of the fractal nature of the organization of depositional sequences proposed by Schlager (2004), which is the expression of the ratio between sediment supply and accommodation at different scales.

Discussion Sequence hierarchy in the Sarvak Formation Five orders of depositional sequences have been distinguished in the studied mid-Cretaceous succession. The Kazhdumi and Sarvak Formations of Albian to Turonian age, covering a time span of approximately 23 Ma, form one large-scale, first-order sequence, bounded by two major disconformities: the Upper Aptian disconformity and the intra-Turonian disconformity. These are recognized at the scale of the Arabian Plate and are the bounding surfaces of the Wasia Group (Harris et al. 1984; Pascoe et al. 1995; Sharland et al. 2001; van Buchem et al. 2002a, b). The disconformities are the expression of major tectonic reorganizations at the scale of the Arabian Plate that took place during the late Aptian (change in direction of plate movements; e.g. Al-Fares et al. 1998) and the late Turonian–Coniacian (compressional regime, obduction along eastern plate margin; Glennie et al. 1990; Ricou 1971). The stratigraphic organization within this largescale sequence can be subdivided in two secondorder sequences, either by using the stratigraphic discontinuity at the top of Sequence I (early Cenomanian, Lower Member of the Sarvak Formation) that testifies to exposure and incision of the platform, or the change in sedimentary system that occurs at the base of Sequence I, when the Albian, muddy, Orbitolina-dominated ramp systems were replaced by the rudist-barrier rimmed platforms. In this instance, the first option has been chosen. The first second-order sequence consists of the Khazdumi formation and the lower member of the Sarvak Formation (Sequence I; Figs 5 & 12). It is composed of at least four third-order sequences with an estimated duration of 2–5 Ma. The first three third-order sequences have a low-angle ramp architecture, with a dominantly muddy sediment composition dominated by benthic foraminifera (mostly oribitolinids and trocholinids; Fig. 5). The last third-order sequence (Sequence I of the Sarvak Formation) is characterized by a platform to intra-shelf basin topography, with rudistdominated barriers, that developed in response to the latest Albian/early Cenomanian transgression (Haq et al. 1988; Hardenbol et al. 1998). This interval corresponds to the MFS K110 of Sharland et al. 2001 and Simmons et al. 2007. Part of the ‘Kazhdumi’ basin was filled in by this system. The second second-order sequence consists of the middle and upper members of the Sarvak Formation, corresponding to three third-order sequences (Sequences II to IV, Figs 5 & 12), with an estimated average duration in the order of 1.5–3 Ma. Here again, the first third-order sequence (Sequence II)

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Fig. 12. Chronostratigraphic scheme of the Bangestan Group in the Padena and Bibi Seydan area (High Zagros). The time scale is from Hardenbol et al. (1998) A. Z., Assemblage Zone as defined by Wynd (1965); MFS, Maximum Flooding Surface.

is characterized by a more tabular, low angle architecture, whereas during the following two third-order sequences (Sequences III and IV) a distinct platform to intra-shelf basin topography developed, in response to the late Cenomanian –early Turonian eustatic sea-level rise (Haq et al. 1988; MFS K140 of Sharland et al. 2001). In detail, at the base of Sequence III, a very flat depositional system remained (Fig. 9). The intra-shelf basin to platform architecture developed during the late Cenomanian (Sequene III) and Turonian (Sequence IV), even though the two third-order transgressive phases were separated by a sea-level drop causing an exposure surface during lastest Cenomanian times. Following the model presented here, the low angle ramps developed during periods of low accommodation when, under the right conditions, the carbonate sedimentation kept up with sea-level rise, and the platform to basin topography developed in periods of high rate of accommodation creation (fast sea-level rise), when the carbonate production was not homogeneous on the platform, causing a differentiation in highs and lows (intra-shelf basins). The evolution of the depositional profile that is observed in this pattern of alternating low-angle, ramp type systems and higher relief platform to intra-shelf basin systems at the second-order scale, is similar to the one reported for third- and fourthorder sequences. This similarity strongly suggests

the fractal character of the organization of sequences in this context and this type of sedimentary systems. The detailed study of the third-order sequences showed these are composed of high frequency fourth- and fifth-order sequences, more or less well expressed depending on the amplitude and frequency of the variations in the ratio accommodation/carbonate production (A/S). Generally speaking, the high frequency cycles are better expressed in the transgressive systems tracts of the third-order sequences. This is particularly clear in the lower Turonian transgressive deposits of Sequence IV (Fig. 11), where an increase of the amplitude of relative sea-level fluctuations is observed.

Correlation with the Natih Formation in Oman The sequence stratigraphic organization of the firstorder Kazhdumi-Sarvak sequence is very similar to the one defined for the Wasia Group in the Oman Mountains by van Buchem et al. (1996, 2002a), Gre´laud et al. (2006) and Homewood et al. (2008) at approximately 1000 km to the SE (Fig. 2a). The main difference with Iran is the input of significant amounts of fine siliciclastic material in the Oman sequences, recorded as Orbitolinid-rich green shales in an inner platform environment during the third-order transgressions (Fig. 13).

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Fig. 13. Stratigraphic correlations between the Sarvak Formation (Zagros, Iran; this study) and the Natih Formation (Oman; after van Buchem et al. 2002a) based on composite synthetic sections. The studied outcrops of these two late Albian–early Turonian Formations are more than 1000 km apart. The remarkable similarities in depositional models and the third-order sequence succession, suggests that eustatic sea-level variation was the main controlling factor of this carbonate system on the Arabian Plate. (1) Dariyan carbonate platform (Aptian); (2) mixed argillaceous-carbonate ramp; (3) Orbitolina carbonate ramp; (4) intra-shelf basin facies association; (5) outer ramp/base of clinoforms facies association; (6) bioclastic platform margin facies association; (7) back-barrier and rudist rich open lagoon facies association; (8) inner platform facies association; (9) karstified chaotic unit; (10) regressive facies sequence; and (11) transgressive facies sequence. SB, sequence boundary; IS1 & 2, platform incisions (see Gre´laud et al. 2006, 2010).

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This difference is due to the relative proximity of the exposed Arabian shield towards the SSE in Saudi Arabia (Fig. 2a). The Albian deposits in Oman are mostly composed of green shales and thin carbonate beds, both rich in Orbitolinids (Nahr Umr Formation, e.g. Immenhauser et al. 1999). These are overlain by the carbonates of the late Albian – early Turonian Natih Formation which have a sequence stratigraphic architecture (sequences, geometries and thicknesses) very similar to the Sarvak Formation in Iran (Fig. 13). Sequence I in Oman, which corresponds to Sequence I in the Iranian study, is 140 m thick and of late Albian– early Cenomanian age. At its base, there are two argillaceous-carbonate units that have a ramp-type architecture (Natih F and G members) and which are overlain by an intra-shelf basin– carbonate platform system (Natih E; Fig. 13). The top of this sequence is a major sub-aerial exposure surface with evidence for platform top incisions resulting from a 20 m or greater fall in relative sea-level (van Buchem et al. 1996, 2002a; Gre´laud 2005; Gre´laud et al. 2006; Homewood et al. 2008). Sequence II, which is 50 m thick and of mid-Cenomanian age, consists of an argillaceouscarbonate lower part (Natih D), and a carbonatedominated upper part (Natih C) which both have a low angle ramp architecture (Fig. 13). In detail, the depositional geometries in this sequence II attest low amplitude intraplate deformations at this time that may result in the relative low accommodation rate and the relative high terrigenous influx. The sequence is bounded at the top by a sub-aerial exposure surface with preserved evidence of pedogenesis. Sequence III, which is 120 m thick and of late Cenomanian and Turonian age, also starts with an argillaceous-carbonate unit with a low-angle ramp architecture (lower part of the Natih B member), which is followed by an intra-shelf basin creation in the late Cenomanian (upper part of Natih B; Fig. 13). This basin is partly filled by a succession of prograding clinoforms of late Cenomanian and Turonian age respectively (see figures Razin & Gre´laud in Homewood et al. 2008). The Turonian progradation is interpreted as a tectonically-forced regression, caused by the first deformations of the compressional tectonics leading to the obduction of the Samail ophiolitic complex. The geometries and the thickness of the sequences recognized in the outcrops of the Natih Formation (van Buchem et al. 1996, 2002a), confirmed and further refined by seismic analysis and iteration of this with outcrop studies (Droste & van Steenwinkel 2004; Gre´laud 2005; Gre´laud et al. 2006; Droste 2010), are very similar to those observed in the sequences of the Sarvak Formation (Fig. 13): (1) growth of the clinoforms and the time-equivalent development of the intra-shelf

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basins during the periods of strong aggradation in sequences I and III; and (2) during sequence II, which is characterized by half the sedimentation rate (and accommodation), no relief develops, and a ramp type architecture exists throughout the sequence. A difference between the two systems is the angle of the clinoforms, which is generally lower in the Natih Formation than in the studied outcrops of the Sarvak Formation, whereas the amplitude, and thus the bathymetry of the intra-shelf basins, is comparable (40 –70 m). The similarities between the sequences in the Natih Formation and Sarvak Formation show that the dynamics of these carbonate systems, and thus the factors controlling them, are homogeneous at the scale of 1000s of kilometres. It is concluded that eustatic sea-level variations are the dominant controlling factor for this particular carbonate system on the Arabian Plate. Small, local, variations in subsidence are easily compensated for by the carbonate system, such as in Sequence II in the Padena outcrop but also in Oman. Only major tectonic deformations, for example during the Turonian in Oman, show an overprint on the stratigraphic expression of the eustatic sea-level fluctuations. The small and local influx of terrigenous material at the beginning of the transgressive phases does not have a major influence on the dynamics of the Natih carbonate system, compared to the purely carbonate dominated Sarvak system. The volume of carbonates produced seems to be controlled by the accommodation, and thus in this case by the eustatic variations in sea-level. It follows the rise of sea-level, except at certain times, when the rise is too fast, causing the development of intra-shelf basins in the areas where carbonate production was unable to compensate for the creation of accommodation. A significant difference that has been observed in the intra-shelf basinal facies between Iran and Oman is the concentration of marine organic matter. This is relatively high in Oman (4% in mature state at outcrop, and up to 15% in immature state in the subsurface; van Buchem et al. 2002a), whereas the organic matter content in the Iranian outcrops is less than 0.3%. It has been demonstrated that the preservation of organic matter in these intra-shelf basins varies considerably (van Buchem et al. 2005), which may go some way to explain this difference.

Depositional sequence model and dynamics of carbonate systems The quality of the outcrops of the Sarvak Formation in the Zagros enables the development of a reference model for the geometrical and facies evolution that

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can be expected in this type of carbonate systems based on the changing ratio of the carbonate production (S) and the accommodation (A) (Fig. 14). The key element in this model is the evolution of the angle of the depositional profile, as controlled by the variations in sedimentation rate on the platform, which are a function of A/S. During periods when the rate of accommodation creation is balanced, the carbonate production tracks the accommodation, and the depositional profile remains (sub) horizontal. In this case, the carbonate system forms a low-angle ramp architecture and aggrades. This is the case during the early transgressive phases (early TST). Above a certain threshold rate of accommodation (in the order of 30–50 m/Ma in this case study) the carbonate factory is not able to fill up the created space, causing a heterogeneous aggradation style on the platform (Fig. 14). A, under these circumstances, is larger than S. Only locally does the carbonate production manage to follow sea-level rise, maintaining a shallow-water environment in a rapidly aggrading setting, whereas sedimentation in other

areas lags behind causing an intra-shelf basinal environment to develop. The creation of these submarine palaeotopographies leads to the development of clinoforms through a differential aggradation rate. These are the late transgressive deposits (late TST). This happens during the major eustatic transgressions of the latest Albian (Sequence I) and the late Cenomanian–early Turonian (Sequences III and IV). However, the lower accommodation rates observed during Sequence II (about half that observed in the other sequences), did not lead to the creation of an intra-shelf basin, neither in Iran nor in Oman. In this system, dominantly controlled by eustatic sea-level variations, the rate of accommodation is never high enough to cause the total drowning of the platform (sensu Schlager 1989). When carbonate production becomes greater than the accommodation, the system starts to fill in the intra-shelf basin with prograding clinoforms varying locally from sigmoidal to oblique (highstand systems tract). In case there is no differential subsidence, and the intra-shelf basin is filled to spill, the result is an almost isopachous sequence,

Fig. 14. Regional depositional model for one third-order sequence in a platform– intra-shelf basin carbonate system (see explanation in the text).

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which has a very complicated internal architecture (Fig. 14). This stratigraphic complexity has been described in Oman (van Buchem et al. 2002a; Droste & Van Steenwinkel 2004; Gre´laud et al. 2006; Homewood et al. 2008), and is clearly documented in seismic work presented by Droste (2010). The evolution of the depositional profile of the carbonate system is accompanied by the evolution of several parameters that have an influence on the dynamics of the system and the carbonate production. (1) The increase in the angle of the depositional profile favours the transition of a low-energy, muddy system to a high-energy, grainy, bioclastic system. This evolution can be interpreted as the result of an increase in the wave energy due to increased water depth and fetch, possibly in combination with an increase of the influence of tidal currents through the larger water volumes that are moved over the shelf. (2) The morphology and the faunal assemblage of the facies model change, with notably the creation of a grainy platform barrier, and the rudist-rich back barrier deposits. (3) The nature of carbonate production changes from initially very fine grained (carbonate mud) of bacterial and algal origin with abundant benthic foraminifera (orbitolinids, prealveolinids, miliolids, etc.) in the ramp setting, to the more granular bioclastic-dominated facies (of mollusc and in particular rudists) dominating the platform margin setting, with a continuing carbonate mud production in the internal platform. Ooliths are absent along the high energy margins of these intra-shelf basins, but are abundantly present in timeequivalent ocean margin deposits (Hillgartner et al. 2003; Razin et al. 2005). (4) The volume of carbonate sediment produced also changes with the variations in relative sea-level. A relatively low production is observed during the early transgression, when the rate of accommodation creation is slow. With increasing rate of sea-level rise, the carbonate production increases, locally to the extent that it outpaces accommodation, leading to prograding geometries before the point of maximum accommodation creation has been reached (maximum flooding surface). During highstand the carbonate production remains high, and with reduced accommodation, the platform is forced to prograde. With gradual exposure of the platform the carbonate producing surface becomes smaller and smaller, to the extent that finally only a thin shelf margin wedge is left. This general evolution is a clear example of the feedback loop that links the carbonate factory, the depositional profile, the sedimentary processes, the ecosystem and accommodation in particular (Homewood 1996). Early transgressive phases are characterized by a relative high volume of nutrients

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and a return of pioneer communities with high organic matter productivity. The depositional profiles are flat or very slightly inclined resulting in wave energy attenuation and mud-rich sedimentation. Late transgressive to regressive phases coincide with low nutrient availability and promote mature or climax communities with high skeletal carbonate production (rudists, stromatoporids, corals, molluscs, etc.). Inclined depositional profiles favour high energy conditions and bioclastic sedimentation along aggrading to prograding clinoforms. In addition, the depositional model has been based on observations made at the third-order scale, but a similar pattern can be recognized at the second-order as well as at the fourth-order scale. The evolution of the system is controlled by variations of the ratio A/S through the feedback loop, but this seems to be independent of the frequency of the variations, suggesting a fractal nature (e.g. Schlager 2004). The depositional model proposed for the midCretaceous carbonate system of the Arabian Plate seems not to be specific for rudist-barrier rimmed platform to intra-shelf basins. A very similar dynamic system has been observed in fourthorder sequences of oolitic dominated depositional sequences in the Lower Cretaceous platform margin of Northern Oman (Lebec 2004). The oolithic barriers of the Habshan Formation (Hauterivian to lower Barremian) are formed during late transgressive stage, when the angle of the depositional profile increased, and the sediment changed from dominantly muddy, to dominantly oolitic (packstones to grainstones). Similar observations have been made in the high-frequency depositional cycles of the late Toarcian to Aalenian oolitic ramp systems in the Moroccan Atlas Mountains (Amelago Formation) (Pierre 2006; Pierre et al. 2010). Although the depositional model predicts the creation of intra-shelf basins during the main phases of transgression, it does not predict their exact location. The studies on the Cenomanian/ Turonian in Iran and Oman show that there was no significant differential subsidence. The mechanism proposed here is similar to the one proposed by van Buchem et al. (2002b) for the creation of the Aptian age Bab intra-shelf basin. A small topographic difference is considered to be sufficient to trigger the differential sedimentation rates during increased rate of relative sea-level rise. This small topographic relief (in the order of a few metres) may have been created in different ways, such as a relict palaeotopography of the previous depositional sequence, local small differences in compaction, small scale tectonic activity related to intra-plate stress regimes or Precambrian to Early Cambrian Hormuz salt movements.

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Outcomes of the stratigraphic correlations The sequence model developed in this paper has several implications for stratigraphic correlations and interpretations of carbonate successions: (1)

(2)

Open marine organisms such as ammonites, brachiopods, planktonic foraminifera and nannofossils have been observed in the shallow-water argillaceous carbonate facies of the early transgressions (early TST) at the base of depositional sequences in the intraplatform domain. Examples are the Natih Formation in Oman (van Buchem et al. 2002a), the Barremian– Aptian Kharaib and Shuaiba Formations in Oman (e.g. Pittet et al. 2002; van Buchem et al. 2002b; Granier et al. 2004), the Jurassic Sahtan Group in Oman (Rousseau et al. 2005), and the lower Jurassic Calcaires d’Amelago in Morocco (Pierre 2006). If environmental interpretations are only based on the presence of these fossils, these units may be mistakenly interpreted as maximum flooding surfaces, at the base of shallowing-up successions, an interpretation that can be contested on facies and geometrical arguments (Fig. 15). The penetration of ‘deeper-water’ organisms into this shallowwater setting, should be attributed to the extremely flat morphology and low energy of the depositional system in the early transgressive phase, before the development of a high energy barrier that would hinder such widespread distribution. Figure 15 demonstrates the implications this difference in facies interpretation can have on the sequence stratigraphic interpretation. The Bibi Seydan example illustrates the complication that variations in sediment supply may cause for the sequence stratigraphic interpretation. The platform margin started to prograde (decrease of A/S) while sea-level still continued to rise, as could be deduced from the depositional geometries (Fig. 16). This can only be explained as a relative increase in sediment production. Following the definition, the MFS is positioned when the system starts to prograde, in this case before the moment of maximum water depth was reached. In a system with continuous sediment supply, such as many siliciclastic systems, the sedimentary response would be in phase with the relative variations in sealevel, and step back, as long as sea-level would rise (Fig. 16). The position of the maximum flooding surface would thus be placed at different positions, even though the absolute sea-level fluctuation would be the same.

Fig. 15. Two distinctive interpretations of a typical Cretaceous inner platform succession: the ‘classic’ interpretation, whereby the argillaceous mudstones (locally with planktonic fauna) are interpreted as maximum flooding surface; and the interpretation proposed in this study, whereby a distinction is made between an early, slow transgression over a flat shelf (hence the deeper-water fauna in a relatively shallow-water setting; see also discussion), and a late, faster transgression, creating accommodation on the platform, and leading to the local development of intra-shelf basins.

(3)

(4)

The importance of volumetric partitioning between the transgressive phases characterized by the aggradation of the internal platform, and the regressive phases when most of the sediment accumulated along the platform margins has been demonstrated for depositional sequences at different scales. This is a fundamental rule to take into account from the point of view of stratigraphic correlations between the platform and the basin. One of the fundamental differences between carbonate systems and siliciclastic systems is not only that the carbonate production takes place in situ, but also that it varies during the course of a depositional sequence. It increases during the transgressive phase and diminishes strongly when the platform becomes gradually exposed. This limits the amount of backstepping in carbonate systems, due to the rapid

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Fig. 16. Timing of platform progradation in response to sediment supply evolution and accommodation cycle. (a) Variable sediment supply: early increase in carbonate production causes the system to prograde before ‘absolute’ maximum water-depth is reached. The MFS of the A/S cycle occurs earlier (T4) than the maximum accommodation point (T6). This situation commonly occurs in carbonate systems. (b) Constant sediment supply: the MFS of the A/S cycle is in phase with the maximum accommodation point (T5). This situation commonly occurs in siliciclastic systems.

(5)

increase in production, but it also limits the amount of forced regressive progradation, due to the rapidly reducing supply of sediment. In addition, the link between increased carbonate production and an increase in the depositional angle of clinoforms during transgression, explains the development of gravitycontrolled resedimentation processes during platform aggradation (Eberli 1991; Razin et al. 2006). This is another difference with siliciclastic systems, where turbidite deposits are mostly deposited during periods of sealevel fall (Vail et al. 1991; Kolla et al. 1995; Mutti 1992). This carbonate depositional model provides elements for the interpretation and reconstruction of grainy carbonate reservoir bodies (Figs 14 & 15). Complex bioclastic and oolitic sedimentary bodies develop progressively during a depositional sequence, and have a maximum extent, in both a platform and basinal direction, at the end of the transgression (backstepping barrier), and during highstand (progradation of the barrier). Both, a diagenetic overprint, related to possible sub-aerial exposure phases, and argillaceous muddy, locally dolomitized carbonate facies of the early transgression are potential low permeable heterogeneities occurring in between the grainy reservoir units.

Conclusions A high resolution sequence stratigraphic model has been constructed for the mid-Cretaceous Sarvak Formation in the High Zagros region of SW Iran, which was located at the eastern side of the Arabian Plate, close to the NeoTethys ocean margin. The exceptional outcrop quality, displaying detailed geometrical and facies patterns in the transition zone from carbonate platform to intra-shelf basin, has allowed distinction between the relative contribution of carbonate sediment supply (S) and accommodation (A) to the architecture of the sedimentary system. A depositional reference model for the Cenomanian/Turonian carbonate systems of the Arabian Plate is proposed based on these outcrops. The main conclusions are: (1) Hierarchical organization of the sedimentary system Four orders of sequences have been defined: two second-order sequences, the first one incorporating the Albian Kazhdumi Formation and the lower part of the Sarvak Formation (first third-order sequence), and the second one covering the middle and upper part of the Sarvak Formation (second to the fourth third-order sequences). The second-order sequences last 10 –15 Ma, are bounded by major exposure surfaces and have lowstand wedges onlapping the disconformity surface. The four third-order sequences making up the Sarvak Formation last between 1.5–3 Ma, and are

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50–150 m thick. These are in turn composed of fourth-order and fifth-order sequences that form the stratigraphic building blocks of this carbonate system. (2) Depositional systems Two main depositional systems alternate in this sequence framework: (a) a shallow-water platform system, with a lowangle ramp geometry, characterized by muddy carbonate sediment, with an admixture of benthic foraminifera and some rudist debris; and (b) a platform to intra-shelf-basin system characterized by mixed muddy/grainy platform top sediments, a grainy rudist-rich, platform margin with low and high angle clinoforms, and a more restricted, organic-matter rich intra-shelf basinal facies in the depression with an estimated maximum water-depth of 80 m. The first type developed from an almost vertical, homogeneous aggradation, with carbonate production keeping up with creation of accommodation creation. This was a typical situation for the early TST of the third-order sequences and the secondorder sequences. The second type of depositional system developed during the main transgressive phases of the late Albian/earliest Cenomanian, and the late Cenomanian/earliest Turonian, when the carbonate system at the scale of the Arabian Plate was not capable of keeping up with relative sea-level rise. The differential aggradation that followed gave rise to the creation of grainy margins with clinoforms that dipped into intra-shelf basinal depressions, and gradually deepened with progressive aggradation. The complete or partial infill of these intra-shelf basins was subsequently achieved by the progradation of the grainy clinoforms when carbonate production exceeded the creation of accommodation in the shallow-water areas. The process that triggered the change from a flat ramp system to an intra-shelf type system was probably the increase in rate of sea-level. This particular step is identified in this study by the distinction between the early (slow) transgression, and the late (fast) transgression. This distinction is an important refinement of the carbonate depositional models. (3) Accommodation Accommodation is thus identified as the dominant factor controlling the morphology of the depositional profile, and, as such, the driving motor behind the dynamics of this type of carbonate system. Through its control on the dip angle of the depositional profile, it has a major influence on: (a) the hydrodynamics of the system, low angle, low energy v. higher angle, higher energy; (b) the type of carbonate sediment, muddy with benthic foraminifera, of bacterial-algal origin in the low-angle system v. grainy, bioclastic, rudist-debris rich barrier system in the high-angle system; and (c) the volume of carbonate sediment

produced, carbonate production increases with gradual creation of grainy barrier along the platform margin. The relative evolution of these different parameters also has an influence on the geometry, the distribution of the sedimentary bodies, and in particular on those with a reservoir potential, during the course of a depositional sequence (feedback loop). The correlations established between the Sarvak Formation and the Natih Formation, at more than 1000 km away in the Oman Mountains, indicate that eustatic sea-level variations were the controlling parameter on this carbonate sedimentary system. In this study, no evidence for tectonic disturbance of this stratigraphic pattern has been observed, other than the Turonian uplift in the Oman Mountains, and probably in the Zagros Mountains. (4) Carbonates versus siliciclastics The interaction between the relative variations in sea-level and the carbonate production, as documented in this study, explains a number of key differences with siliciclastic systems: (a) the response of the carbonate system to fast accommodation increase is the switch from a flat ramp type depositional system to an intra-shelf to platform type system. This involves changes in the depositional geometries, faunal composition and sediment composition. This critical change is emphasized in this paper by the distinction between an early TST (characterized by a relatively slow rise in relative sea-level) and a late TST (characterized by an increase in the rate of relative sea-level). This change in rate of sea-level rise causes an immediate backstepping in siliciclastic systems, but no significant change in the depositional geometries or depositional facies; (b) during sub-aerial exposure of the carbonate platform, the carbonate factory is restrained to a thin fringing wedge, causing a diminution of the carbonate production. This limits the extent of the forced regressive wedges, and thus the infill of the intra-shelf basins. In the case of siliciclastic systems, where sediment supply is more continuous, there is no such limitation. (5) Fractal nature Although reasonably complex, the depositional model for the evolution of the geometries and dynamics of this carbonate system in an A/S cycle is applicable at different orders of sequences (second, third and fourth), and appears to be applicable in different types of carbonate producers (rudists, ooids) and at different ages. This study thus supports the hypothesis of a fractal nature of sedimentary sequences.

CRETACEOUS PLATFORM MARGIN IN SW IRAN This study is part of the PhD thesis of FT which was financed by the Petroleum University of Technology (N.I.O.C), Tehran, Iran, and hosted by the Institut Franc¸ais du Pe´trole in Rueil-Malmaison, France, and the University of Bordeaux III, Institut EGID, France. The authors would like to acknowledge the logistic support for the fieldwork campaigns provided by the Exploration Directorate of the National Iranian Oil Company, and in particular Mr. Mohaddes, Mr. Zadeh-Mohammadi, Mr. Ahmadnia, Dr. Syooki and Dr. Baghbani. Field assistant support by NIOC colleagues M. Kavosi, A. Bakhshi, M. Jalali and M. Khosravi is greatfully acknowledged. Thin sections were prepared by the Palaeontology Department of the NIOC. Dr. Bahrami and Prof. Schroeder have assisted in the palaeontological identifications. Drawing assistance by Y. Monteon is gratefully acknowledged. Constructive comments from reviewers M. Simmons and P. Homewood have helped to improve the manuscript.

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Regional stratigraphic architecture and reservoir types of the Oligo-Miocene deposits in the Dezful Embayment (Asmari and Pabdeh Formations) SW Iran F. S. P. VAN BUCHEM1,2*, T. L. ALLAN3, G. V. LAURSEN4, M. LOTFPOUR5,6, A. MOALLEMI5, S. MONIBI5, H. MOTIEI7, N. A. H. PICKARD8, A. R. TAHMASBI9, V. VEDRENNE1,10 & B. VINCENT1,11 1

IFP, Geology-Geochemistry division, 1 – 4 Avenue de Bois Pre´au, 92852 Rueil-Malmaison, France

2

Present address: Maersk Oil Qatar AS, P.O. Box 22.050, Doha, Qatar

3

CSIRO, 11 Julius Avenue, Riverside Corporate Park, North Ryde 1670, NSW, Australia 4

Statoil ASA, Grenseveien 21, NO-4035 Stavanger, Norway

5

RIPI, West Boulevard, Azadi Sports Complex, Tehran, 14665-1998, Iran

6

Present address: KEPS, 50 Emami Street, North Sohrevardi Avenue, Tehran, 15516-13816, Iran 7

NIOC R&D, No.22, Negar Street, Vali-Asr Avenue, Vanak Square, Tehran, 19698-13771, Iran

8

Cambridge Carbonates Ltd/CarbRes AB; Gullvivevagen 5, 756-55, Uppsala, Sweden

9

NIOC Exploration Directorate, Geology Department, 1st Deadend, Seoul Street, NE Sheikh Bahaei Square, Tehran, 19395-6669, Iran 10

Present address: ENSPM, IFP School, 228– 232, Avenue Napoleon Bonaparte, 92852 Rueil-Malmaison, France

11

Present address: Cambridge Carbonates Ltd; 14 rue du Mont, 52320 Marbeville, France *Corresponding author (e-mail: [email protected]) Abstract: A regional sequence stratigraphic model is proposed for the Oligo-Miocene Asmari and Pabdeh Formations in the Dezful Embayment of SW Iran. The model is based on both new detailed sedimentological observations in outcrops, core and well logs, and an improved high-resolution chronostratigraphic framework constrained by Sr isotope stratigraphy and biostratigraphy. A better understanding of the stratigraphic architecture distinguishes four, geographically separated types of Asmari reservoirs. Three Oligocene sequences (of Rupelian, early Chattian and late Chattian age) and three Miocene sequences (of early Aquitanian, late Aquitanian and early Burdigalian age) have been distinguished, representing a period of 15.4 Ma. The stratigraphic architecture of these sequences is primarily controlled by glacio-eustatic sea-level fluctuations, which determined the distribution of carbonates, sandstones and anhydrites in this sedimentary system. Tectonic control became important in the Burdigalian with a regional tilt down towards the NE. The lithological heterogeneity, the complex geometries, and both early and late diagenetic alterations are the basis for a classification of four main stratigraphic reference types for the Asmari Reservoirs: Type 1, sandstone dominated; Type 2, mixed carbonate-siliciclastic; Type 3, mixed carbonate-anhydrite; and Type 4, carbonate dominated. The sequence stratigraphic model predicts how and when these types change laterally from one to another.

The Oligo-Miocene deposits in the Dezful Embayment of SW Iran are a complex sedimentary system with a large variation in lithologies (sandstones,

marls, carbonates and anhydrites) and depositional geometries, and, as a consequence, a diachroniety of the formation boundaries. The moderate to

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 219–263. DOI: 10.1144/SP329.10 0305-8719/10/$15.00 # The Geological Society of London 2010.

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shallow-water carbonates and sandstones of this system (Asmari Formation) are important hydrocarbon reservoirs, which have been producing oil since the early twentieth century. With fields reaching the latter stages of their producing lifecycle, there is now an increasing interest in enhanced oil recovery methods to optimise the production. One of the initial stages in this process is obtaining an improved understanding of the reservoir heterogeneities, which are controlled by four factors: lithological variation, depositional geometries, diagenesis and structural deformation (fractures, faults, etc.). In this paper we will focus on the basin scale sequence stratigraphic architecture of this system. Comprehensive studies of the depositional history of the Oligo-Miocene deposits in the Dezful Embayment have been published in the 1960s [with lithostratigraphic work by Thomas (1950), James & Wynd (1965); internal company reports by Adams & Bourgeois (1967) and Adams (1969); and microfacies work by Sampo (1969)]. Although these studies provide information relevant for the regional setting of the Oligo-Miocene palaeogeography, they predate recent improvements in the resolution of biostratigraphic and chemostratigraphic dating, and the application of seismic and sequence stratigraphy. Recent publications on the Asmari Formation are few, and of a more local nature, such as the outcrop work by, for example, Seyrafian & Hamedani (1998), Seyrafian (2000) and Vaziri-Moghaddam et al. (2006), and the subsurface work on diagenesis by Aqrawi et al. (2006) and the Sr isotope work by Ehrenberg et al. (2007) and Mossadegh et al. (2009). These papers propose preliminary sequence stratigraphic interpretations of the Asmari Formation based on a dataset of a limited number of wells or outcrops covering only part of the Dezful Embayment. Of general interest for this study are the publications on time-equivalent Formations in Northern Iraq by van Bellen (1956, 1959), Dunnington (1958) and van Bellen et al. (1959), and more recently by Al-Hashimi and Amer (1985, 1986), Goff et al. (1995) and Aqrawi et al. (2010). The main purpose of this paper is to present a basin scale, sequence stratigraphic model, which explains and predicts the regional variations in depositional facies and reservoir facies. A key aspect of this study was the establishment of (high resolution) timelines across the different depositional facies and lithologies. Recent work by Ehrenberg et al. (2007) in the Dezful Embayment showed the potential of Sr isotopes for use in dating these Cenozoic rocks. This methodology has also been applied systematically in this study and has improved the age control significantly, to a resolution of approximately 0.5 Ma. This technique was also used to recalibrate the commonly used index fossils, and propose a revision of the

age assignment of the biostratigraphic zonation scheme. Seismic scale outcrops of the Asmari and Pabdeh Formations can be studied in the Dezful Embayment close to the oilfields that produce from the Asmari reservoirs. This unique situation enables the outcrop analogues to be used in both the context of regional stratigraphic reconstructions, and also in the study of reservoir scale heterogeneities. In addition, a good regional understanding of the Oligo-Miocene stratigraphic architecture is also required to ensure the selection of appropriate outcrop analogues for the different types of Asmari reservoirs observed in the fields. The adapted methodological approach was to compare and contrast the varying data sources, including sedimentological outcrop observations, strontium isotope dating, biostratigraphy, well log and seismic interpretations, to propose the most internally consistent sequence stratigraphic framework of the studied rock succession. The results presented in this paper have been acquired as part of a larger project that covered the stratigraphic, diagenetic and structural aspects of the Oligo-Miocene deposits in the Dezful Embayment, preliminary results of which have been reported by Daniel & Azzizadeh (2006), Daniel et al. (2008), Laursen et al. (2006, 2009), van Buchem et al. (2006, 2009), Vincent et al. (2006) and Nader et al. (2008). The subject of this paper is the regional sequence stratigraphic architecture. First, the sedimentological facies are presented, then the Sr isotope stratigraphy and revised biostratigraphy, followed by the depositional sequences. Finally the implications of the proposed integrated stratigraphic model for the regional stratigraphy and Asmari reservoirs are discussed.

Geological setting The study area is located SW of the Zagros Mountains, which is a young, Mio–Pliocene, fold-thrust belt located along the eastern margin of the Arabian Plate (Fig. 1a). The Zagros is the result of a multi-phased collision between the Arabian plate, the former southern margin of the Neo-Tethys Ocean, and the Central Iran Microplates. The structural map is characterized by the main NW –SE trend of the Zagros thrust belt, within which two major thrusts can be distinguished, the Mountain Front Fault and the High Zagros Fault (Fig. 1b). These major thrusts are not rectilinear, and are offset by major transverse lineaments such as the Izeh-Hendijan and Kazerun Fault Systems (Fig. 1b). Structural information derived from seismicity within the Zagros belt proves the existence of these strike-slip faults in the basement underlying the folded cover (Jackson 1984; Ni 1986; Berberian 1995; Talebian 2002; Talebian & Jackson 2004).

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Fig. 1. (a) Geographical map with positions of studied wells and outcrops. The two transects discussed in this paper are labelled Transect 1 (in red) and Transect 2 (in blue), and are presented in Figure 17. Other transects used for this study are indicated in black. (b) Topographic map of part of the Zagros Mountains showing the main structural elements. The study area is indicated by the black rectangle. ZFTB, Zagros Front Belt.

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They correspond to linear uplifts along pre-existing basement trends (Motiei 1993). These major tectonic lineaments define several well known morphotectonic units (Falcon 1974; Haynes & McQuillan 1974; Favre 1975; Berberian 1995). The study area is mainly situated in the Dezful Embayment and Izeh Zone, but includes also outcrop sections in the High Zagros (Fig. 1). The Izeh Zone is part of the Zagros simply folded belt, and is characterized by numerous outcrops of the Asmari Formation. The Dezful Embayment structural zone is characterized by a low elevation and few outcrops of the Asmari Formation, and it contains most of the Iranian oil fields (Cenozoic and Mesozoic). The present day morphology of the Zagros Mountains is the result of a long geological history,

including the Tethyan rifting phase in the Permian – Triassic; the passive margin phase, with seafloor spreading towards the NE, in the Jurassic–Early Cretaceous; the first compressive phase with subduction to the NE and obduction in the Late Cretaceous; and, finally, the second compressive phase with collision during the Neogene (e.g. James & Wynd 1965; Ricou 1974; Motiei 1993; Sherkati & Letouzey 2004). The Late Cretaceous compressive phase is of particular importance since at that time the NW– SE trending foreland basin was created that controlled overall sedimentation on the Arabian Plate in the Palaeogene (Fig. 2). Paleocene and, in particular, the Eocene platforms first developed around the margins of this basin, prograding in towards the centre. The Oligocene/early Miocene

Fig. 2. Paleogeographical maps. (a) Middle Eocene paleogeography and plate reconstruction (from C. R. Scotese 2001). Inset shows the palaeogeography at the late Eocene, with shallow-water carbonate platform deposits of the Jahrum Formation in the south and shallow-water carbonates and possibly exposure in the north (based on Adams & Bourgeois 1967; Sampo 1969; Goff et al. 1995; Aqrawi et al. 2009 and the present study). The NW–SE orientation of the basin was inherited from the Late Cretaceous foreland basin (Goff et al. 1995). (b) Middle Miocene paleogeography and plate reconstruction (from C. R. Scotese 2000). Note the advanced closure of the NeoTethys.

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Fig. 3. Lithostratigraphic scheme. Schematic presentation of the main lithostratigraphic units encountered in the study area along a southeast to northwest transect. Note the significant lateral variation in lithology. The Middle Anhydrite corresponds in Lurestan to the Kalhur Member.

Asmari platforms represent the final stage of this long-lived progradational carbonate platform system, which came to an end with the closure of the NeoTethys Ocean and subsequent collision phase during the late Miocene (Fig. 2) (Homke 2004). The studied stratigraphic interval includes the following lithostratigraphic units (Fig. 3): (a) the shallow-water carbonates of the Eocene Jahrum Formation; (b) the mixed, shallow-water carbonates and sandstones of the Asmari Formation; and (c) the deeper water marls of the Pabdeh Formation. The latter formations interfinger with two subaqueous anhydrite intervals, the Basal and Middle Anhydrites. In addition, in the southwestern part of the basin a sandstone unit, the Ahwaz Sandstone Member is distinguished within the Asmari Formation. Oligo-Miocene sedimentation was initially controlled by the palaeo-topography of the preceding Eocene Jahrum platform. This platform surrounded a large embayment that was open towards the NeoTethys and extended into northern Iraq (Fig. 2). Following a major eustatic sea-level drop at the close of the Eocene (e.g. Abreu et al. 1998; Sharland et al. 2001), shallow-water sedimentation downstepped in this basin, while the surrounding carbonate platforms were exposed. The Asmari Formation was deposited on this first order sequence boundary (base of tectono-stratigraphic megasequence 11 of Sharland et al. 2001).

Methods and materials This basin scale study integrated sequence and seismic stratigraphy, sedimentology, biostratigraphy and chemostratigraphy to produce a regional correlation. Examples in the literature of sequence stratigraphic methodology and definition of principles, can be found in for example, van Wagoner et al. (1990), Loucks & Sarg (1993), Read et al. (1995), Homewood & Eberli (2000), and Sharland et al. (2001). The subdivision of depositional sequences into five orders, which fall into a general time framework, is followed here (Vail et al. 1991; Haq et al. 1987). The orders of sequences of relevance for this study are: second order (3–50 Ma), third order (0.5–3 Ma), fourth order, also referred to as high frequency cycles, para-sequences or genetic sequences (0.5–0.08 Ma), and fifth order (0.08 –0.02 Ma). Their definition is based on the notion of geological time, and they are the result of tectonic, tectono/eustatic, eustatic, or glacioeustatic mechanisms (Vail et al. 1991). The definition of sequences in this sequence stratigraphic model is taken from van Wagoner et al. (1990) and Vail et al. (1991), who define a sequence as limited by a sequence boundary (SB) at the top and the base, with a maximum flooding surface (MFS) separating the transgressive (TST) and highstand (HST) systems tracts. A sequence boundary with evidence

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for exposure is a Type I boundary, one characterized by a shallowing up trend is a Type II boundary. The following symbolic convention was used to assign depositional sequences: (a) upward pointing blue triangle corresponds to an Accommodation (A) over Sediment supply (S) ratio that is larger than 1 (A/S . 1), there is more space created than can be filled by the sedimentation, resulting in a backstepping (landward stepping) stacking pattern of depositional cycles, or a transgressive trend; and (b) a downward pointing red triangle corresponds to an A/S , 1, there is more sediment produced or supplied than space created, resulting in a forward stepping (seaward stepping) stacking pattern of depositional cycles, or a regressive trend. The methodological approach consists of the following steps: (1) The detailed macroscopic sedimentological and palaeontological description and sampling of outcrop sections and core. This included observations on discontinuity surfaces, stratal patterns, and, where possible, depositional geometries. (2) Complementary laboratory analyses, including semi-quantitative microscopic analysis of the micro-facies, strontium isotope analyses, and biostratigraphic analyses. (3) The definition and environmental interpretation of the sedimentary facies. (4) One dimensional analysis of the depositional sequences and the establishment of a sequence hierarchy (i.e. stacking pattern analysis) (5) The construction of regional log and outcrop section correlations, incorporating the available time control (biostratigraphy and chemostratigraphy) and geometrical control (outcrops, seismic). The cross reference of the sequence definition throughout the basin led to the final definition of the sequence stratigraphic framework. (6) The construction of palaeogeographical maps and isopach maps using the established sequence framework. (7) Definition of the sequence stratigraphic model and chronostratigraphic scheme.

Materials This study used a combined outcrop and subsurface dataset of both existing data, such as wells, palaeologs, and seismic, and newly acquired data, including new outcrop sections and newly described cores. New analyses were performed on the latter. The 51 control points cover an area of approximately 300  300 km (Fig. 1b), and comprise: † 10 new surface sections, measured at the 1 : 100 scale (Katoola, Izeh, Chidan, Kuh-e-Asmari,

Kuh-e-Bangestan, Kuh-e-Khaviz, Kuh-eKhami, Kuh-e-Razi, Tang-e-Gurguda, Jahadabad), covering a total of 4564 m of stratigraphic section. All sections are located on the geographical map (Fig. 1b), and GSM coordinates are provided in Table 1. † core material of 10 wells described at the 1 : 50 scale, covering a total of 2305 m. † 31 wells with existing palaeologs (a semiquantitative microfacies analysis performed on cuttings). † A limited number of vintage seismic lines. Systematic sampling for thin section preparation was used in the outcrop logging, with a sample frequency of one sample every 2– 3 m. In the cored wells the microfacies analyses were performed on samples taken every 30 cm, in un-cored wells thin sections were made from the cuttings, generating a continuous microfacies log along the wells. Although occasional errors introduced in the cutting samples due to caving and imprecise positioning of the samples cannot be excluded, this method is considered to provide a very accurate dataset, which formed one of the key data sources of this study. It was thus possible to interpret both outcrop and well data in terms of depositional environment within the entire study area. Biostratigraphic analyses were carried out on 10 outcrop sections and 10 cored wells by A.-R. Tahmasebi (NIOC Exploration Directorate), S. Monibi, Z. Karimi (RIPI), and N. Moradi (independent consultant), advised by G. V. Laursen (Statoil) and J. P. Masse (Total). Semi-quantitative microfacies analysis (Palaeologs) for all new outcrop sections (1175 thin sections), and the newly described core (9042 thin sections) were completed. Sr isotope dating by measurement of the 87 Sr/86Sr ratio has been carried out on 256 samples in 9 surface sections and 6 wells. The method is based on the known variation in seawater 87 Sr/86Sr through the Phanerozoic (Burke et al. 1982; McArthur & Howarth 2004). The majority of the analyses were made on whole rock subsamples, identified as the most effective method based on the limestone types and project scale. Samples were screened in thin sections for diagenetic effects. Sampling and analyses were performed in the laboratories of the CSIRO Petroleum Division in Sydney. For each sample a ‘whole-rock’ powder was taken with a dental drill, sampling sparry or micritic matrix and bioclasts, but with a bias to avoiding coarsely crystalline calcite or dolomite. Fossil isolates were either drilled or separated with a pneumatic ‘wiggler’ drill. Powders and fossil fragments were dissolved in 1 M acetic acid to minimize extraction of Sr from clays and other terriginous

Outcrop sections

Thickness

Thin sections

Biostat.

Sr samples( ): measurem.

Easting

Northing

X (Easting)

Y (Northing)

2000000 1964160 1973900 1932720 1987050 2051900 2066220 2075870 2094500

1099000 1116980 1073440 1088450 1010870 972523 944543 946160 948000

432575.86 397771.37 405062.31 364751.89 414698.24 477325.81 490054.89 499781.80 518488.40

3516574.45 3536553.43 3492498.46 3509804.82 3429247.30 3387311.51 3358565.55 3359637.74 3360426.29

6

Outcrop dataset Katoola Izeh Chidan Kuh-e-Asmari Kuh-e-Bangestan Kuh-e-Khami Tang-e-Gurguda Jahadabad Kuh-e-Razi

651 m 1358 m 480 m 353 m 137 m 375 m 515 m 310 m 385 m

114 257 104 171 50 87 173 136 92

Total

4564 m

1175

1

2

3

4

5

– – –

– – –

– –

– –

– –

Well dataset Well number Lithology Biostratigraphy Chemostratigraphy Core

        

15 (15) 21 (24) 17 (18) 14 (15) 16 (16) 23 (27) 49 (54) 6 (6) 19 (20)

UTM Datum 186 WGS84 Zone N39

Lambert Iranian

180 7

– –

8

9

– –

10

11

– –

12

13

– –

STRATIGRAPHIC ARCHITECTURE AND RESERVOIR TYPES OF THE ASMARI FORMATION

Table 1. List of GSM coordinates of the outcrop sections

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material (e.g. DePaolo et al. 1983), and strontium extraction followed standard methods of ion exchange. This whole-rock method has been used in similar scale projects in the Papuan Basin (e.g. Whitford et al. 1996; Eisenberg et al. 1996; Allan et al. 2000). Isotopic compositions were measured on a VG 354 thermal ionization mass spectrometer. External precision of +0.000015 (2 standard errors of the mean) is assumed from measurement of the NIST 987 (or NBS 987) standard over the period of analysis. Strontium isotope ages were determined from the McArthur & Howarth (2004) look-up table Version 4: 08/03 for the seawater strontium curve, calibrated to the GTS2004 time-scale of Gradstein et al. (2004). 87Sr/86Sr ratios reported here are normalized to NIST 987 87Sr/86Sr ¼ 0.710235, and are 0.000013 less than the McArthur & Howarth (2004) values.

Sedimentology The sedimentary facies of the Asmari and Pabdeh Formations are organized in the main lithologies

of these formations: carbonates and marls, siliciclastics and anhydrite. Carbonates, marls and anhydrites are all well represented in the studied outcrops and core material. The siliciclastic facies has, however, only very locally been found in outcrop, and a limited amount of core of the siliciclastic facies has been described. A certain bias towards detail in the carbonate and marl facies is thus partly based on the limited availability of siliciclastic facies for study. The interpretation of the depositional environment is a key step in the definition of depositional sequences. In this mixed sedimentary system, a facies substitution diagram is used to compare the environmental interpretation of the facies defined for the different lithological classes (Fig. 4).

Carbonates and marls The carbonate facies classification is based on the Dunham texture and the macro- and micro-fossil content. The carbonate and marl facies have been grouped in facies associations that represent five main depositional environments, ranging from the intertidal to the basinal setting (Table 2). The

Fig. 4. Facies substitution diagram. This diagram shows the spatial distribution of different sedimentary facies encountered in the study area along a virtual depositional profile, with in the lower part the siliciclastic facies and in the upper part the carbonate dominated facies.

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Table 2. Carbonate and marl facies types, facies associations and depositional environments Carbonate and marl facies 1. Subaerial exposure 1a: Conglomeratic dissolution breccia in carbonate facies 2. Inner ramp/platform (intertidal to upper subtidal) – restricted benthic fauna 2a: Fenestral wackestone and dolomitized mudstone; locally with anhydrite nodules 2b: Oolitic, peloidal packstone/grainstone; with dm- to m-scale x-bedding 2c: Miliolid-rich, wackestone/grainstone; dominantly benthic foraminifera with additional faunal elements (ostracods, green algae) 2d: Mixed small benthic foraminifera wackestone to packstone 3. Open platform top/mid ramp(subtidal) – diverse benthic fauna From proximal to distal: 3a: Skeletal wackestone/packstone; dominated by echinoids, bivalves, oysters, bryozoan, red algae and locally microbialite encrustation; dm-scale x-bedding and channels 3b: Coral dominated floatstone/boundstone; both platy and large (dm to m scale) domal morphologies are observed. Associated fauna consists of red algae, echinoids, benthic foraminifera 3c: Red algal dominated rudstone to boundstone; this facies has two occurrences, either as layers consisting of rhodolitic rudstones (rhodoliths diameters up to 5 cm), or as a red algal boundstone. 3d: Large benthic foraminifera wackestone/packstone; typical fauna contains Nummulites, Operculina, Heterostegina, Amphistegina and small forms of Eulepidina, associated with red algae, echinoderms, bivalves. 4. Platform slope/Outer ramp – dominantly benthic with some planktonic elements 4a: Wackestone to packstone with a mixed benthic/planktonic fauna; typical faunal elements include Globigerinids, bryozoan, ditrupa, large foraminifera (e.g. large forms of several centimetres in diameter of Lepidocyclinas, Eulepidinas), echinoderms and red algae. Locally (transported) miliolids can be mixed in. 4b: Favreina packstones; a coprolite facies typical for the margins of the intrashelf basin. 5. Basin/hemi-pelagic – dominantly planktonic fauna locally with some benthic elements 5a: Pelagic foraminifera dominated mudstone to wackestone.

environmental interpretation of the facies is based on: (a) faunal assemblages; (b) depositional geometries observed in outcrop and seismic; (c) lithological information; and (d) the existing international literature on microfacies interpretation (e.g. Thomas 1952; Sampo 1969; Buxton & Pedley 1989; Franseen et al. 1996; Seyrafian & Hamedani 1998, 2003; Aqrawi et al. 2006; VaziriMoghaddam et al. 2006), and internal NIOC/ OSCO reports. A summary of the environmental interpretation of most common micro- and macrofauna is shown in Figure 5, where the faunal elements have been plotted along a virtual depositional profile. It should be noted that some of these faunal elements were not coeval, due to the biological evolution, and that different faunal assemblages are thus characteristic of different time periods, as will be discussed below. These assemblages are: Assemblage A: dominated by Nummulites, and occurring in the Eocene and earliest Oligocene; Assemblage B: dominated by a mixed NummulitesEulepidina fauna, occurring in the early Oligocene (Rupelian); Assemblage C: dominated by Eulepidina, but without Nummulites, occurring in the late Oligocene (Chattian);

Assemblage D: dominated by Favreina (a crustacean coprolite facies), occurring in the Aquitanian due to particular environmental conditions (see below). The depositional geometries observed at outcrops are generally characterized by horizontal bedding of decimetre- to metre-scale beds. Two exceptions are: coral buildups and large scale clinoforms. Coral buildup geometries have been observed in two locations: in Kuh-e-Razi, where a several kilometres long escarpment displays a 60 m thick accumulation of corals with a depositional relief in the order of 30–40 m, and which passes laterally in deeper water deposits (Fig. 6b); and in Kuh-e-Khaviz, where coral buildups of a smaller dimension have been observed, with a topographic relief of 3 –6 m (Fig. 6a). Spectacular clinoforms, with a topographic relief in the order of 90 m, have been observed along a several kilometres long escarpment in the southern flank of Kuh-e-Mish (Tang-e-Gurguda; Fig. 7b, c). It shows that at least during part of its history, the Asmari carbonates formed (locally) steep platform to basin topography. Since there is no evidence that this was the common depositional geometry, a general classification of a ramp model is followed in the environmental characterization of the facies.

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Fig. 5. Depositional profile and distribution of marine carbonate producers. Carbonate producing organisms encountered in this study are presented along a virtual depositional profile with their estimated living position and area of distribution. This figure is based on the literature cited in the text and our own findings.

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Fig. 6. Coral buildup facies. (a) Kuh-e-Khaviz; small relief coral buildups are overlaying offshore sandstones enriched in Eulepidinas. (b) Kuh-e-Razi; high relief coral buildups observed along a several kilometres long escarpment displaying clear onlap geometries.

Siliciclastics

(2)

Sandstones have only been observed in three outcrop localities, and in the core material of a limited number of wells. As a result the facies description and interpretation is very succinct. Three main depositional environments can be distinguished: (1)

Coastal plain to terrestrial environment— Quartz sandstones and clays without fossils, locally with root traces (palaeosoils) have been observed in core material. They occur interbedded at the decimetre- to metre-scale with platform top carbonates. This facies association is interpreted as a coastal plain to fluvial environment.

(3)

Distal deltaic and subtidal environment—2a: Metre-thick, massive, sandstone interbedded between carbonate platform top deposits (Katoola outcrop section). These facies are interpreted as shallow marine sandstones, probably deposited in the subtidal domain of the platform top environment. 2b: Cleaning up trends in GR logs, which are interpreted as coarsening and shallowing up trends from basinal marls, containing a pelagic fauna, to fossil-poor sandstones, as supported by cutting analysis (e.g. Well-5). This facies association is interpreted as prograding, deltaic sand bodies. Offshore marine to basinal setting—3a: Wellsorted, medium-grained quartz sandstones, deposited as a several metre thick sheet sand in between open marine carbonates

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Fig. 7. Tang-e-Gurguda outcrop section. (a) Composite log of sedimentological observations, strontium isotope measurements and biostratigraphy. The age assignment is based on the Sr isotope data. (b) Overview of the Pabdeh Formation and the lower part of the Asmari Formation seen from the northern flank of the Tang-e-Gurguda gorge in the Mish Anticline. Note the NNE prograding clinoforms in the lower part of the Asmari Formation. (c) Detail of prograding

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Fig. 7. (Continued ) clinoforms in the lower part of the Asmari Fm. The clinoforms have a topographic relief in the order of 90 m, and are continuously exposed for several kilometres. This geometrical observation is highly significant for two reasons, firstly it provides key information on the depositional processes and related geometries of the Asmari limestones, and secondly it provides evidence for the magnitude of sea-level rise in Sequence II.

(Kuh-e-Khami outcrop section). These sands are interpreted as reworked offshore deposits. 3b: Yellow-red weathered, 2–3 m thick finegrained sandstone beds, locally with abundant Eulepidinas, interbedded with marls and passing upward into carbonates (northeastern

side of Kuh-e-Khaviz; Fig. 6a). These sands are interpreted as reworked offshore deposits, probably more distal than facies 3a. The source of most of the sandstones was located in the west, where the exposed Arabian Shield shed

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Fig. 8. Katoola outcrop section. (a) Composite log of sedimentological observations, strontium isotope measurements and biostratigraphy. The age assignment is based on the Sr isotope data. (b) Overview of the Late Cretaceous and

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Fig. 8. (Continued) Cenozoic stratigraphic succession at the Katoola road section. (c) Overview of the upper part of the Asmari Formation in the Katoola road section. The base of the Gachsaran is marked by a several metres thick anhydrite.

large amounts of clastics onto the eastern Arabian plate (e.g. Sharland et al. 2001). In the Katoola section, however, the analysis of the mineralogical composition of the sands (dominance of pyroxene) suggests a provenance of basic, probably ocean floor, origin. Further detailed analyses are needed, but preliminary results suggest that these sands were probably not sourced from the Arabian Shield in the west, but rather from a source in the east where obducted radiolites and ophiolites were eroded.

Anhydrites Anhydrites have been observed at three stratigraphic levels in the Asmari Formation: (1) The first level is the ‘Basal Anhydrite’, which has been observed in two locations. In Kuh-eAsmari, a 4 m thick anhydrite bed is encased in marls containing a pelagic fauna (Fig. 9b). Directly below the anhydrite, wavy-bedded

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Fig. 9. (a) Kuh-e-Asmari section: composite log of sedimentological observations, strontium isotope measurements and biostratigraphy. The age assignment is based on the Sr isotope data. (b) Kuh-e-Asmari section: overview of the lower part of the Kuh-e-Asmari section with the Pabdeh marls, including the Basal Anhydrite, and the lower part of the Asmari Formation limestones. Insert 1 shows a detail of stromatolitic features found at the base of the Basal Anhydrite. Insert 2 shows the basal anhydrite interbedded between the deep water marls of the Pabdeh Formation. (c) Kuh-e-Bangestan section, overview of the lower part of the Asmari Formation including the Basal Anhydrite interbedded in the limestones of the Asmari Formation. Insert shows detail of the anhydrite bed.

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Fig. 9. (Continued ).

and slightly domed stromatolites are found. In Kuh-e-Bangestan, the ‘Basal Anhydrite’ is 6 m thick, and interfingers with shallow-water carbonates (Fig. 9c ), where they have a sharp contact with the carbonates below and above. (2) The second level is the ‘Middle Anhydrite’, which has only been observed in the wells, where it reaches a maximum thickness of 70 m

and consists of alternating anhydrite beds (8–12 m thick) and dolomite beds of similar thickness (Well-13; Fig. 10). (3) The third level, present both in outcrops and in the subsurface, comprises a several metres thick anhydrite bed that overlies the Asmari Formation and forms the base of the Gachsaran Formation (e.g. Katoola outcrop section;

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Fig. 8c). In the Chidan outcrop section, a wavy to domal, thinly layered, stromatolitic bed is found just below the anhydrites. Considering the thickness and lateral continuity of the Basal and Middle Anhydrite deposits, it is assumed that they formed subaqueously in an isolated high salinity basin. The exact water depth of these deposits is difficult to estimate, but may be in the order of 10 or 15 m, based on the basin morphology. The presence of the stromatolitic deposits at the base of the anhydrites, observed in Kuh-e-Asmari and in Chidan, suggests either a subaqueous origin of these particular stromatolites, or a sea-level fall to near exposure. Considering the regional context, the former interpretation would be favoured for the stromatolites at the base of the basal anhydrite, and the latter for the stromatolites at the base of the Gachsaran Formation. A full understanding of the significance of these stromatolite deposits would, however, require further study.

Lithostratigraphy Three outcrop sections and three well locations have been chosen to illustrate the geographical variability in the lithological composition of the Oligocene –Miocene sedimentation pattern. These are presented in Figures 7 to 12, and show examples of the composite datasets that have been collected.

cross-bedding. At 450 m a well developed karst surface is observed.

Katoola outcrop section The Katoola section is located at the northeastern margin of the basin. The Asmari Formation is carbonate dominated, has a thickness of 630 m (Fig. 8a), and overlies a thick succession of Paleocene/earliest Oligocene marls of the Pabdeh Formation (Fig. 8b). The top is marked by an anhydrite bed of several metres thickness, marking the base of the Gachsaran Formation (Fig. 8c). The Asmari Formation ranges here in age from early Rupelian to Burdigalian, and is one of the thickest and longest ranging carbonate successions in the study area. The lower part of the Asmari is represented by a massive carbonate cliff (0–170 m) of Rupelian age, that shows an overall shallowing-up trend from basinal marls of the Pabdeh Formation at the base, to (distal) slope carbonates with Nummulites and Eulepidina, and a miliolid-rich facies at the top. From 170–510 m the environment is an open platform top setting, with a well developed bedding pattern, consisting of an alternation of limestones and marl-rich intervals (Fig. 8c). At 210 m, a several metres thick siliciclastic unit has been observed, and at 240 m is a well developed karst surface. The top 120 m, from 510–630 m, consists of pelagic marls of Burdigalian age.

Tang-e-Gurguda outcrop section The Tang-e-Gurguda section is located at the southeastern margin of the basin. The Asmari Formation is carbonate dominated, reaches a total thickness of 450 m (Fig. 7a), and overlies marls of the Pabdeh Formation (Fig. 7b). The top of the section is probably the contact with the Gachsaran Formation, but no anhydrite beds were preserved in this location. The Asmari Formation ranges here in age from the early Rupelian to the Burdigalian. In the lower part of the Asmari Formation (75– 130 m), a shallowing-up trend to a mid ramp environment is observed, expressed in a facies evolution from basinal marls of the Pabdeh Formation to a Eulepidina-rich facies of Assemblage C (Fig. 7a), although Thomas (1951) reported Nummulites for this part of the section, in which case it would classify as the Nummulites–Eulepidina fauna of Assemblage B. This is followed by a hundred metre thick interval characterized by the presence of large scale clinoforms, with a relief in the order of 90 m, indicating an important deepening (Fig. 7c). From 230 –510 m, the environment is a shallow-water platform top setting, with coral packstones and boundstones and miliolid-dominated facies, locally with metre- to decimetre-scale

Kuh-e-Asmari outcrop section The Asmari Formation was defined by James & Wynd (1965) at Kuh-e-Asmari. This location is in the basin centre, and the Asmari Formation reaches here a total thickness of approximately 350 m (Fig. 9a). At the base of the outcrop, late Oligocene deep water marls of the Pabdeh Formation are exposed, including a 5 m thick anhydrite bed (Fig. 9b). For comparison, an illustration of a similar anhydrite bed observed in the Kuh-eBangestan section is shown in Figure 9c. The outcrop in Kuh-e-Asmari, stops close to the top of the Asmari Formation, but no anhydrite bed of the Gachsaran Formation has been found. The age of the Asmari Formation ranges from early Aquitanian to Burdigalian at this location. The deposition of the anhydrite bed was a shortlived event, after which deeper water marl deposition continued. The transition to shallow-water limestones occurs in a massive cliff of 90 m (Fig. 9b), where a complex facies succession is observed from marls with a mixed benthic and pelagic fauna, to red-algae and bryozoan-rich beds, to echinoderm wackestones with oysters, to a Favreina-dominated facies, eventually followed

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Fig. 10. Well-13. Composite log of well logs, sedimentological observations and biostratigraphy. The age assignment is based on Sr isotope calibrated index fossils. Note the presence of both the Basal Anhydrite, interbedded in basinal marls, and the Middle Anhydrite, occurring at the change to shallow-water carbonates.

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Fig. 11. Well-5. Composite log of well logs, sedimentological observations and biostratigraphy. The age assignment is based on Sr isotope calibrated index fossils. Note the presence of sandstones of the marine Ahwaz Sandstone Member in the lower part of the Asmari Formation.

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Fig. 12. Well-2. Composite log of well logs and sedimentological observations and biostratigraphy. The age assignment of the Jahrum Formation is based on index fossils, the age assignment of the Asmari Formation is based on regional correlations (see text). The Eocene carbonate platform of the Jahrum Formation is here overlain by a relatively thin Asmari Formation, that consists mostly of sandstones of the Ahwaz Sandstone Member.

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by miliolid wackestones at the top. From 90 –350 m the Asmari formation has been deposited in somewhat restricted to open carbonate platform top conditions (Fig. 9a). The bedding pattern is well developed with decimetre- to metre-scale beds.

Well-13 Well-13 is also located in the basin centre. The Asmari Formation consists here of carbonates, anhydrite beds and marls, reaches a thickness of 340 m, and overlies a 300 m thick succession of marls of the Pabdeh Formation of Paleocene to early Oligocene age (Fig. 10). The top of the Asmari Formation is marked by the anhydrite deposits of the base of the Gachsaran Formation. The age of the Asmari Formation ranges from the base of the Aquitanian to the early Burdigalian. The Lower part of the Asmari Formation consists of two anhydrite beds, the ‘Basal Anhydrite’ of about 12 m thickness, and the ‘Middle Anhydrite’ of about 50 m thickness, which are separated by a 90 m thick interval of basinal marls (Fig. 10). From 450 –650 m the environment is interpreted as a carbonate platform top setting.

Well-5 Well-5 is located at the southwestern margin of the basin. The Asmari Formation consist here of sandstones (Ahwaz sandstone member) and carbonates, reaches a thickness of 435 m, and overlies 310 m of basinal marls of the Pabdeh Formation of Paleocene/early Oligocene age (Fig. 11). The top of the Asmari Formation is marked by anhydrite deposits of the base of the Gachsaran Formation. The age of the Asmari Formation ranges from the Chattian to the early Burdigalian. The lower 200 m of the Asmari Formation (350–550 m; Fig. 11) is predominantly sandy (Ahwaz Sandstone Member). The sandstones have been interpreted, based on the log signature, stratigraphic position and the correlation with nearby wells, as offshore sands, possibly distal delta lobes, and represent an overall shallowing-up trend. The overlying carbonates (135 m) contain a fauna indicating a restricted platform top environment.

Well-2 Well-2 is located in a proximal location at the southwestern margin of the basin. The Asmari Formation has a thickness of 200 m, and is dominantly sandy. It lies directly on top of the Eocene platform deposits of the Jahrum Formation (Fig. 12). The top of the formation is marked by an anhydrite bed that marks the base of the Gachsaran Formation. The Asmari Formation is here probably mid Chattian to early

Burdigalian in age. Since no microfauna have been found in the sandstones, the age was estimated based on regional correlations. The lower part of the Asmari Formation consists of 170 m of sandstones (Ahwaz Sandstone Member), and the upper part consists of 30 m of carbonate. The sandstones do not contain faunal elements. Some clay interbeds occur. More information is needed to interpret their environment; it is probably very shallow marine to continental.

Biostratigraphy and Sr isotope stratigraphy Although the Asmari Formation is one of the most prolific oil producing sequences in the world, relatively little is known of its detailed stratigraphic palaeontology. Thomas (1950, 1952) outlined the fundamental biostratigraphy of the Asmari Formation based on large benthic foraminifera. This is the only published record of the biostratigraphy before James & Wynd (1965) formally defined the formation. In addition, the traditional Iranian biostratigraphy is based on unpublished reports (Wynd 1965; Adams & Bourgeois 1967) as described in Seyrafian & Hamedani (1998). These reports were written in a period where the Aquitanian was a subject of debate (e.g. Eames et al. 1962; Berggren 1963; Drooger 1964). Many workers regarded the Aquitanian as being wholly or partly Oligocene (see discussion in Eames et al. 1962). In older publications, such as those of Henson (1950), Smout & Eames (1958), and the reports of Wynd (1965) and Adams & Bourgeois (1967), microfossils ascribed to the ‘Aquitanian’ may in fact refer to sediments that would be assigned at the present day to the Chattian of the late Oligocene. Aside from these problems with chronostratigraphic definition of the foraminiferabased zones, facies restriction of biostratigraphic markers limited basin-wide application. A solution to this dating problem would be to calibrate the ranges of index fossils to an alternative dating method, which is independent of depositional environments, such as strontium isotope stratigraphy. Chemostratigraphically defined time lines can be followed from the platform margin settings, where sequences have a clear sedimentological expression, to both platform top and basin environments, where the sedimentological and faunal information is often not clear enough to identify sequences. With the rapid increase in the seawater 87Sr/86Sr ratio during the Oligocene to Middle Miocene, strontium isotope stratigraphy can potentially resolve accurate age profiles through the section, to identify the presence of stratigraphic gaps, and to quantify the sedimentation rates. Following encouraging results

STRATIGRAPHIC ARCHITECTURE AND RESERVOIR TYPES OF THE ASMARI FORMATION

obtained with Sr isotope dating of the Asmari Formation by Ehrenberg et al. (2007), it was decided to systematically sample key outcrop sections and cored wells for Sr isotope measurements, and compare these dating results with the distribution of the index fossils.

Sr isotope stratigraphy A total of 256 samples, distributed over 9 surface sections and 6 wells, were analysed for Sr isotopes. A selection of this dataset is presented in the outcrop sections of Figures 7–9, and in the compilation graphs of Figures 13 and 14. The observed 87 Sr/86Sr ratios vary between 0.707889 and 0.708561, corresponding to ages of between 32.3 Ma and 18.2 Ma (early Oligocene to early Miocene) applying the strontium seawater curve of McArthur & Howarth (2004). In both the well and surface sections the 87Sr/86Sr ratios systematically increase up section, while the biostratigraphic ages broadly validate the general accuracy of strontium ages derived from these ratios. It is therefore assumed that the measured 87Sr/86Sr trends in each section have largely preserved the global record of strontium evolution in oceanic seawater. Sources of error in the accuracy of Sr ages include uncertainty in the seawater Sr curve, measurement uncertainty (analytical error), and any isotopic heterogeneity in the material analysed. Quantifying the error associated with a Sr age is therefore difficult. Through the Oligocene to early Miocene (to c. 19.5 Ma), the error on the Sr ages associated with both measurement uncertainty and ‘curve’ error (lookup tables of McArthur & Howarth 2004) is c. +0.3–0.4 Ma. By about 19.0 Ma this combined error is c. +0.2 Ma. However, these values do not account for any heterogeneity in Sr-isotopic composition due to diagenesis, contamination from non-marine Sr, or unquantifiable error in the seawater Sr curve itself. With respect to the latter, a plateau occurs in the long-term 87Sr/86Sr evolution at the Chattian/ Aquitanian boundary, as shown by Cahuzac et al. (1997). As a consequence, this critical boundary is better constrained by Sr-isotope stratigraphy in some sections, but is difficult to accurately define using the method. Replicate whole-rock (n ¼ 5) and multicomponent (fossil/matrix n ¼ 18) analyses conducted in this study indicate that isotopic heterogeneity within the Asmari limestones is similar to or slightly greater than the analytical uncertainty, equating to age differences of between 0.2 and 1.2 Ma within any particular sample. Fossil analyses in 6 samples yielded measurably higher 87Sr/86Sr ratios (younger ‘ages’) than the host matrix. In the context of Sr age trends in the sections

241

concerned, it is uncertain if the (younger) fossil Sr ages are more accurate. However, diagenesis (dolomitization and fabric destructive non-dolomitic recrystallisation) has clearly resulted in stratigraphically restricted anomalously old whole-rock ‘ages’ in the Chattian at Tang-e Gurguda (at 300 and 350 m; Fig. 7). Ehrenberg et al. (2007) reported a similar general fossil-matrix age difference of around 0.3 Ma, with younger fossil ages. They interpreted the older matrix ages as possibly due to diagenetic exchange of Sr from older limestones. In the Chidan, Katoola, Tang-e Gurguda, Kuh-eKhami, Kuh-e-Bangestan, Kuh-e-Razi sections and possibly Well-9 (Fig. 14), there is a small age reversal of c. 1.0 to 2.0 Ma in dolomitized and nondolomitized limestones in the Burdigalian. Analysis of Gachsaran Formation bedded anhydrite from Katoola and Chidan also yielded slightly older strontium ages than underlying limestones (Fig. 14). These age reversals may be a systematic and regional effect, and possibly due to seawater strontium in this region deviating in composition from coeval oceanic seawater. However, this is not resolvable on the available strontium dataset. The strontium ages for bedded anhydrite from the lowest Asmari at Kuh-e-Asmari (Fig. 9) and Kuh-e-Bangestan indicate their isotopic composition is probably close to coeval sea water. In summary, our study indicates that stratigraphically restricted diagenetic noise in the Asmari strontium isotopic record is potentially an impediment to correlation, but much of this noise is likely to be small with respect to analytical or biostratigraphic age error. Furthermore, Sr age anomalies have been identified in the context of good biostratigraphic, sequence stratigraphic and petrographic control. Overall, the dataset is believed to provide a sound basis for regional chronostratigraphic correlation in the Dezful Embayment. The interpretation of the chemostratigraphic curves depends on the sample density and the understanding of the sedimentary succession. When dealing with a constant average sedimentation rate, such as in Kuh-e-Asmari (Fig. 9), an almost linear extrapolation can be used between sample points. However, when there are breaks in the sedimentation, as for instance noticed in Well-9 (Fig. 14), a detailed sampling is essential to identify the position of the condensed horizons. In the Katoola section an uncertainty of this nature exists, where a sparse sampling between 230 and 350 m, has left ambiguity in the interpretation (Fig. 8).

Revised biostratigraphy To investigate further the interpretation of the Chattian –Aquitanian and the reliability of the index fossils, the Sr isotope ratios have been

242 F. S. P. VAN BUCHEM ET AL. Fig. 13. Strontium isotope curves for the stratigraphic sections of Transects 1(a) and 2(b). Curves are plotted against stratigraphic height, which allows identifying changes in sedimentation rate and stratigrahic hiatuses. The positions of the sequence boundaries are indicated with stars in the curves, and connected. They show the high consistency of the Sr isotope based ages of the sequence boundaries.

STRATIGRAPHIC ARCHITECTURE AND RESERVOIR TYPES OF THE ASMARI FORMATION 243

Fig. 14. Sr isotope curves, depositional sequences and systems tracts. The Sr isotope curves allow distinguishing between changes in sedimentation rate along transects, with vertical trends indicating high sedimentation rate, and near horizontal trends indicating slow sedimentation or a hiatus. See Figures 1 and 2 for location map of transects. (a) Transect 1. The high time resolution obtained with Sr isotope dating allows distinguishing in some of the sequences between the TST and HST. In particular sequence 4, where the strong aggradation during the TST, and the basinward shift of sedimentation during the HST is demonstrated. Also note the hiatus observed in Well-9 which represents the HST of sequence 4 and all of sequence 5. (b) Transect 2. Note the reduced thickness of sequence 6 as compared to Transect 1, which is due to condensed sedimentation and non-deposition. Also, note the horizontal shift at the amalgamated SB IV and V, marking the stratigraphic hiatus corresponding to sequence 5.

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plotted along with the fossil ranges against the thickness of the outcrop sections and wells (see e.g. in Figs 7 –9). The ranges of the different foraminiferal species are estimated in absolute ages (following the GTS 2004 time-scale of Gradstein et al. 2004) using the Sr ratio derived age interpretation. The interpretation of these results, carried out on 15 outcrops and wells throughout the Dezful Embayment, is summarized as total ranges for the index species in Figure 15. The following conclusions were drawn from the results of this recalibration: (1) revision of the Chattian – Aquitanian boundary, which is now higher in the section (a thinner Aquitanian); (2) demonstration of the environmental dependence of several index fossils, notably Peneroplis evolutus and Austrotrilina howchini; and (3) demonstration of the reliability of index fossils Nummulites spp. for the Rupelian, Archaias hensoni and Miogypsinoides complanatus for the Chattian, and Borelis melo curdica for the Burdigalian (Fig. 16). The revised biostratigraphic zonation scheme illustrated in Figure 15 was introduced by Laursen et al. (2006, 2009), and is described below. The new biozones are described in ascending order, following the stratigraphic development. † The Globigerina spp. – Turborotalia cerroazulensis – Hantkenina Assemblage Zone is defined as the interval containing dominant Globigerina spp. in which the extinction of Hantkenina and Turborotalia cerroazulensis occur. Assemblages like this, with high numbers but low diversity and made up mainly of small globigerinids, usually indicate early Oligocene (especially if they contain Chiloguembelina species, however Chiloguembelina is not observed in our dataset). This age affiliation is confirmed by the Sr-dates. Turborotalia cerroazulensis has been observed in the early Oligocene as well (Pearson 1998). When the assemblage includes Hantkenina spp. the age will be Eocene. † The Nummulites vascus – N. fichteli Assemblage Zone is defined as the total range of the two nominated species of Nummulites. Associated fauna: Operculina complanata, Heterostegina spp., Rotalia viennoti, Eulepidina dilatata, Haplophragmium slingeri, and Ditrupa. Additional species include Eulepidina elephantina, Subterranophyllum thomasi, and Archaias operculiformis. This zone is confined to the Rupelian. † The Lepidocyclina – Operculina – Ditrupa Assemblage Zone has this associated fauna: Planorbulina spp., Heterostegina spp., Eulepidina dilatata, Haplophragmium slingeri, Rotalia viennoti, and algae. This zone ranges from the Rupelian into the Chattian.

† The Archaias asmaricus – A. hensoni – Miogypsinoides complanatus Assemblage Zone is defined by the concurrence of Achaias hensoni, A. asmaricus, Miogypsinoides complanatus and Spiroclypeus blanckenhorni. The zone is restricted to the Chattian. † The Miogypsina – Elphidium sp. 14 – Peneroplis farsenensis Assemblage Zone is defined by the concurrence of Miogypsina spp., Elphidium sp. 14, Peneroplis farsenensis and occasionally Favreina asmaricus. Miogypsina spp. can sometimes occur in large numbers as a ‘flood’. The zone is restricted to the Aquitanian. † Indeterminate Zone. A very fossil poor interval is often observed in between the ‘Miogypsina– Elphidium sp. 14 – Peneroplis farsenensis Assemblage Zone’ and the topmost ‘Borelis melo curdica – B. melo melo Assemblage Zone’. Often the fauna only consists of unidentified Miliolids and Dendritina rangi. Since no identifying species are present, it has been named the ‘Indeterminate Zone’. The zone is mainly associated with the Aquitanian. † The Borelis melo curdica – B. melo melo Assemblage Zone is defined as the total range of Borelis melo curdica within the Asmari Formation. Associated fauna: Dendritina rangi, Meandropsina spp., Spirolina spp., Polymorphinids, Discorbids, small Peneroplids and Peneroplis evolutus, Miliolids and Echinoid debris. The zone is restricted to the Burdigalian.

Sequence stratigraphy The sequence stratigraphic model developed in this paper has been constructed using an integrated dataset including: (1) a high resolution time framework, combining the revised biozonation and Sr-isotope stratigraphy results presented above; (2) the environmental interpretation of depositional facies, based on sedimentological and micropalaeontological observations in outcrop and core; and (3) seismic and outcrop controlled depositional geometries. Apart from well constrained age dating, depositional geometries visible in seismic data or in seismic scale outcrops are theoretically the most important criterion for the correlation of depositional sequences. In this study only a limited number of seismic lines were available to constrain the regional well log correlations. These illustrated in particular the onlap geometries of Oligocene strata against the Eocene palaeotopography. At all studied outcrops horizontal bedding patterns were observed, while at two locations also large scale depositional geometries were found: clinoforms in the Tang-e-Gurguda section (Fig. 7c) and coral build-ups in the

STRATIGRAPHIC ARCHITECTURE AND RESERVOIR TYPES OF THE ASMARI FORMATION 245

Fig. 15. Revised biostratigraphic zonation scheme. Summary of the stratigraphic ranges of Oligo-Miocene benthic foraminifera determined with Sr isotope dating and a proposal for a revised biozonation scheme (see text for further explanations).

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Fig. 16. Comparison of previous biostratigraphic zonation schemes and the one proposed in this study. The main change is subdivision of the previous Aquitanian interval in a much reduced Aquitanian upper part and a Chattian lower part, which implies a repositioning of the Oligo-Miocene boundary.

Kuh-e-Razi section (Fig. 6b). The sedimentological and faunal criteria for sequence definition were particularly efficient in basin margin settings, where a clear environmentally controlled evolution of faunal assemblages and trends in the lithology was observed (e.g. carbonate cleaning up trends, increasing and decreasing sandstone trends, subaqueous anhydrite deposition). In pelagic and platform top domains, however, where environmental changes are reduced and sequence subdivision using sedimentological evidence is often much more difficult, other ways of correlating sequence boundaries were applied, notably Sr isotope dating. The study area is covered by seven cross-sections (Fig. 1b), two of which have been illustrated here (Fig. 17). These two transects include all the outcrop sections and the majority of the cored wells, and are thus best constrained in terms of age control, facies description, and bedding pattern. The sequence and time framework is sufficiently robust to correlate third order sequence boundaries and maximum flooding surfaces from well to well over tens of kilometres. Bed correlation at the metre scale is, however, hazardous over these distances. The fine scale, high resolution correlation is more adapted to field scale studies, where well spacing is in the order of several hundreds of metres to kilometres. This paper focuses at the large scale, regional stratigraphic architecture that will provide a framework for future studies of the finer scale, fourth and fifth order organization of the depositional system. The two cross-sections have been flattened on SB VI, which is the base of the last sequence, and corresponds to the Aquitanian– Burdigalian

boundary (Fig. 17). There are two reasons for this decision, firstly because this surface marks a clear change in the sedimentation style (see below), and secondly it is very well constrained by the appearance of the benthic foraminifera Borelis melo curdica and by Sr isotope dating. The isotope curves for the sections and wells located along Transects 1 and 2 have been plotted in Figures 15 and 16. These illustrate the overall good coherency of the sequence boundary ages (see also Table 3), the gradual younging of the shallow-water carbonates towards the basin centre, and the uniformly high sedimentation rates in the Burdigalian. Six depositional sequences were identified. The duration of the sequences varies, based on the GTS2004 time-scale, from a minimum of 1.2 Ma to a maximum of 5.2 Ma (Table 3), which classifies them as third to very short second order sequences following the time-based classification by Vail et al. (1991). Based on the time control provided by the revised biozonation and strontium isotope stratigraphy and the facies and geometrical control provided by the sequence stratigraphy (SB and MFS surfaces) a chronostratigraphic scheme has been constructed (Fig. 18). This scheme illustrates the presence of hiatuses and the dynamic evolution of the sedimentary system through time. These depositional sequences are described below in stratigraphic order.

The Oligocene sequences The basin morphology in the earliest Oligocene is strongly controlled by the underlying

Fig. 17. Regional sequence stratigraphic correlations and facies distribution. Both transects are flattened on SB VI (base Burdigalian), see Figures 1 and 2 for location map. (a) Transect 1. Note the onlap of Sequences 1 and 2 in the SW against the Eocene Jahrum platform that was exposed during most of the (Lower) Oligocene. Subsequent gradual progradation and infill of the basin occurred from the NE and the SW, with the deposition of the Basal Anhydrite in Sequence 4 (Fig. 9b). In Sequence 5 this basin was filled to spill. A clear lithological difference is apparent in the Asmari Formation with the mixed siliciclastic and carbonate sedimentation along the southwestern margin, the marls, carbonates and anhydrites in the centre of the basin, and the carbonate-dominated sedimentation along the northeastern margin. During sequence 6 a regional tilt occurred towards the NE, creating substantial accommodation represented by a marine transgression (Katoola section; Fig. 8), whereas in the SW sedimentation was strongly condensed or absent. (b) Transect 2. Note the progradation of the carbonate platform along the basin axis towards the NW. Coral-rich margins developed during Sequences 2 and 3 (Fig. 6), and clinoforms during the transgression of Sequence 2 (Tang-e-Gurguda; see Fig. 7b, c). In this transect, both the Basal Anhydrite (Fig. 9b, c) and the Middle Anhydrites (Fig. 10) are present. Note the absence of Sequence 5 in the southeastern sections.

Ehrenberg et al. (2007)

This study Sequence

Surface

2001 Time-scale (Ma)

GTS 2004 (Ma)

Sequence 6

SB VII mfs VI SB VI

18.5 19.6 20.5

18.5 19.4 20.2

Sequence 5

mfs V SB V

20.9 21.8

20.7 21.4

Sequence 4

mfs IV SB IV

22.4 24.2

21.8 23.2

Sequence 3

mfs III SB III

25.1 25.8

24.3 25.1

Sequence 2

mfs II SB II

27.4 28.9

27.2 28.8

Sequence 1

mfs I SB I

32.5 33.9

32.6 34

Duration (Ma)

Sharland et al. (2001/2004)

Sequence

Surface

–

Ng 20 SB Ng 10 mfs –

1996 Time-scale (Ma)

SEPM Tethyan basins 1998 GTS 2004 (Ma)

19 20

1.7 (early Burdigalian)

Bu-20

1.2 (late Aquitanian)

Aq20/Bu-10

1.8 (early Aquitanian)

Aq10

– Ng 10 SB

Ch30

Pg 50 mfs –

24.5

1.9 (late Chattian)

Ru30/Ch-10

Pg 40 mfs –

29

3.7 (early Chattian) 5.2 (latest EoceneRupelian)

–

Pg 30 mfs Pg 30 SB

33 33.5

– – 23

Sequence

GTS 2004 (Ma)

Bur3 SB

18.12

Aq3/Bur1 SB

20.43

Aq2 SB

21.44

Ch4/Aq1 SB

23.03

Ch3 SB

24.84

Ch1/Ru 4 SB

28.45

Pr 4/Ru 1 SB

33.9

23.5 29.1 32.8

STRATIGRAPHIC ARCHITECTURE AND RESERVOIR TYPES OF THE ASMARI FORMATION

Table 3. Ages of sequences and comparison with other studies

247

248 F. S. P. VAN BUCHEM ET AL. Fig. 18. Chronostratigraphic scheme for Transect 1 (see Fig. 1 for location map) showing the sequence stratigraphic organization, with related hiatuses, and an interpreted sea-level curve for this study area. Note the decrease of the amplitude of sea-level fluctuations up section. The studied interval is strongly influenced by the transition from a green-house climate to an ice-house climate which occurred at the Eocene/Oligocene boundary, and which subsequently controlled sedimentation by high amplitude glacio-eustatic sea-level fluctuations (Miller et al. 1991, 2005). The high influx of siliciclastics during the Chattian is probably related to the combination of climatic change and high amplitude sea-level change (see Discussion). During the Burdigalian, at the top of the studied interval, tectonic influence dominated the sedimentation pattern with a strong regional tilt towards the northeast causing a marine transgression in that area.

STRATIGRAPHIC ARCHITECTURE AND RESERVOIR TYPES OF THE ASMARI FORMATION

paleotopography of the Jahrum Formation. At that time, the Nummulites-rich Jahrum platform formed a prominent positive feature in the southwestern part and in the northeastern part of the Dezful Embayment (Fig. 2; Sampo 1969), surrounding a seaway that is known to continue to the NW, into Iraq (Dunnington 1958; van Bellen et al. 1959; Goff et al. 1995), while, towards the SW, a connection with the NeoTethys is presumed. Late Eocene exposure of the southwestern margin of this platform is suggested by the observation of incised channels on the top of this platform on 3D seismic of the Dorquain field (NIOC Exploration information), while ‘Eocene Red Beds’ have been observed at the NE margin (Geological Map; Huber 1969). The age of these proximal sediments is however poorly constrained. Large parts of the Jahrum Formation remained exposed during the early Oligocene, and formed the southwestern margin of the basin, where onlap geometries have been observed (Transect 1, Fig. 17a). From the early Oligocene onwards, shallow-water sedimentation started to occur in the northeastern part of the study area, forming a barrier along the NeoTethys Ocean, and thus creating an isolated basin in the central part of the Dezful Embayment. The Oligocene shallow-water faunal assemblage shows a change in the Rupelian, from a Nummulitesdominated fauna (Assemblage A) to a mixed Nummulites/Eulepidina-dominated fauna (Assemblage B). In the latest Rupelian the Nummulites disappeared, and a Eulepidina-dominated fauna remained (Assemblage C). The disappearance of the Nummulites is an age indicative phenomenon, while the presence or absence of Eulepidina is an environmental indicator (Laursen et al. 2006). The composition of the sediment also changed around the Rupelian –Chattian boundary. During the Rupelian the sedimentary system was carbonate dominated, with marls in the basin, while from the early Chattian onwards significant amounts of sandstones and silts were brought into the basin, which accumulated either close to the source along the southwestern basin margin or were transported to the basin centre (Fig. 17a). Sequence 1 (Rupelian). Sequence boundary I is placed at, or just below the Eocene/Oligocene boundary. In the SW it is identified as a Type I sequence boundary by evidence of non-deposition of the Rupelian sequence (Well-1 and Well-2 in Transect 1; Fig. 17a). In the other sections of the study area this boundary is observed in basinal facies, where it is a Type II sequence boundary, picked at the top of a cleaning-up trend in the gamma-ray log pattern (Figs 10 & 11). This trend is interpreted as an increase in the carbonate

249

content marking maximum regression. The top boundary of this sequence (SB II) is dated as latest Rupelian, just below the Rupelian –Chattian boundary. SB II is in the proximal domain a Type I sequence boundary by absence of the Rupelian sequence. Downslope, on top of the Rupelian shallow-water carbonates it is probably a Type II sequence boundary, capping faunal Assemblages A and B, and locally overlain by (marine ?) sandstones (e.g. Well-3 and Well-4 in Transect 1; Fig. 17a). In the distal domain, this surface is picked on logs above a slight cleaning up trend in the gamma-ray log and based on the faunal content (Figs 10 & 11). In the Katoola section (Fig. 8) there is some ambiguity about the exact positioning of this boundary. Nummulites spp., disappear at 120 m in the section, but a Rupelian age based on Sr isotopes is recorded up to approximately 230 m. Assemblage C, which is assumed to be of Chattian age, occurs, however, between 120 and 170 m, while from a sedimentological point of view the base of several metres of siliciclastic sediments recorded at 210 m may be a suitable candidate for the top of this sequence. As a compromise, 170 m has been chosen as the position of sequence boundary II, with a second option at 120 m (between Assemblages B and C). This section requires more work to further refine this sequence boundary. The general lithology in this sequence consists essentially of limestone in the shallow-water domain, and marls in the deeper water domain. Sands are only very locally recorded in proximal locations in the south. In the basin, a characteristic, regionally expressed shale peak, expressed as a high gamma-ray spike and named the ‘X-marker’ (Figs 10 & 11), occurs just above the basal sequence boundary and is interpreted as a LST. This shale interval is interpreted to have been deposited when the Jahrum carbonate platform became exposed, and the carbonate supply to the basin was at its lowest. This marker is overlain by an interval that is characterized by a flood of small Globigerina spp. (Biozone no. 55 by Wynd 1965) which is interpreted as the basinal expression of the TST and HST of this sequence (Figs 10 & 11). In the more proximal positions a distinction can be made between the southwestern margin, where a clear onlap relationship has been observed of the Rupelian carbonate platform against the exposed Jahrum platform (Wells-3 and 4 in Transect 1; Fig. 17a), and the southeastern and northeastern margins where the Rupelian carbonate platforms evolved during a shallowing up trend that started from late Eocene/early Oligocene deep water marls (Katoola section in Transect 1, Fig. 17a, and the Kuh-e-Razi to Tang-e-Gurguda sections in Transect 2, Fig. 17b). The shallow-water

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fauna consisted in the more proximal areas of Assemblage A (Nummulites dominated, no Eulepidina), while in the more distal position a mixed assemblage of Nummulites and Eulepidina was found (Assemblage B). Sequence 2 (early Chattian). The sequence boundary at the base (SB II) is described above. The top sequence boundary, SB III, is dated as early late Chattian. It is picked in the proximal domain, above the Jahrum platform, in the sandstones below a shale break (Well-2, Fig. 12). The interpretation of this surface is ambiguous in the absence of core and dating. On the basis of regional correlations and resulting geometrical considerations, an exposure, and thus a Type I boundary seems, however, likely in this area (Fig. 17a). In the more distal sections along the southwestern margin, this surface is placed on top of carbonate platform deposits which are generally overlain by sandstones of the following sequence (e.g. Wells-3, 5 and 6 in Transect 1; Fig. 17a). Sr-isotope dating has helped to identify this surface in the wells along the southwestern and southeastern margins, where no evidence for exposure has been observed. In the Katoola section, located in the carbonate dominated northeastern margin, this surface is placed at a karstification surface, providing good evidence for subaerial exposure (240 m; Fig. 8). The lithological composition in this sequence is mixed carbonate/siliciclastic in the SW, and carbonate dominated in the SE and NE (Fig. 17). Lowstand sandstone deposits have been identified in the basin centre, both in wells and in the Kuh-e-Khaviz exposures (Figs 6a & 17). During the transgression, sands continued to be deposited on the platform and in the offshore domain in the southwestern part of the basin (Fig. 17). The log signature shows cleaning up trends, suggesting an architecture of stacked parasequences and/or delta lobes in an overall transgressive setting (e.g. Well-5, Fig. 11). The creation of relief during the transgression is clearly expressed in the Tang-e-Gurguda section, in the eastern part of the study area, where large scale clinoform geometries with a topographic relief of 90 m have been observed prograding towards the NW along a 10 km long escarpment (Fig. 7b, c). During the highstand sedimentation was again more carbonate dominated along the southwestern margin, and characterized by fossil Assemblage C, with Eulepidina, but no Nummulites. The sand influx had either decreased, or sand was stored in the more proximal domains. Along the southeastern margin significant coral buildups developed in the proximal locations of Kuh-e-Razi and Tang-e-Gurguda during transgression. The presence of coral mounds in the HST at the distal position at Kuh-e-Khaviz (Fig. 6a) suggests a

significant fall of sea-level towards the end of this sequence. Sequence 3 (late Chattian). The sequence boundary at the base (SB III) is described above. The top of this sequence, SB IV, is positioned at the OligoMiocene (Chattian–Aquitanian) boundary, and is well constrained with Sr isotope data in both the platform and basinal setting. This boundary marks a major, sea-level drop (see Discussion below), with evidence for subaerial exposure in the surrounding platforms (Type I boundary). The sedimentological expression of this surface is dependent on the palaeogeographical position. In the basin this boundary is picked just below the ‘Basal Anhydrite’, which is also controlled with Sr isotope dating. In the shallow-water carbonate platform deposits the position of this surface is controlled by both sedimentological observations (e.g. karst surface in the Tange-Gurguda section) and Sr isotope dating. In the southwestern part of Transect 1, SB IV is picked just below a laterally continuous sandstone bed, interpreted as the lowstand of the next sequence (Fig. 17a). In the most proximal position of the southwestern part of the area, the sedimentation is sandstone dominated and, here, the boundary has been picked based on the gamma-ray log signature and thickness analogy along the transect. Here incisions and channelling may be expected. Seismic observations have indeed shown channelling in the upper part of the Asmari Formation, but the exact stratigraphic positioning of these is unresolved. The lithological composition of this sequence is mixed siliciclastic/carbonate in the southwestern part of the basin, and carbonate dominated in the northeastern and southeastern part (Fig. 17a). Along the southwestern part of the basin lowstand sandstones occur, followed by transgressive landward stepping sandstones observed in the proximal wells (Wells-2, 3, 4; Fig. 17a). Highstand deposits are carbonate dominated again, except in the most proximal parts in the SW where sands dominated, or no sediments accumulated at all. In the southeastern part of the basin carbonate sedimentation continued. During transgression coral buildups developed in the margin in the Kuh-e-Khami area (Fig. 17b). The platform top environment, established during highstand is here typically characterized by a dominance of miliolids, suggesting a more restricted depositional environment (Fig. 17b). The gross depositional environment of this mixed system is interpreted as one in which transgressive sandstones aggraded on the platform and gradually stepped back during transgression, thereby creating an environmental window for carbonate sedimentation to dominate sedimentation during the highstand.

STRATIGRAPHIC ARCHITECTURE AND RESERVOIR TYPES OF THE ASMARI FORMATION

The early Miocene sequences The Miocene sequences are characterized by two events: (1) eustatic sea-level drops during the Aquitanian, which twice caused the isolation from the NeoTethys and the deposition of submarine anhydrites in its centre (the ‘Basal Anhydrite’ and ‘Middle Anhydrite’ which is time-equivalent to the Kalhur Formation in the NW); and (2) regional tilt of the study area during the Burdigalian, causing uplift and non-deposition in the SW and a shift of the depocentre to the NW where subsidence was high and sediment accumulated along a new margin facing NeoTethys. The faunal assemblage in the Aquitanian was influenced by isolation from the open ocean, which caused an increase in the salinity of the intrashelf basin. The resultant specific environmental conditions led to a restricted, stressed environment; ooids became important allochems in the depositional system as well as Favreina coprolites (Assemblage D). During the Burdigalian, when the intra-shelf basin was in-filled in the study area, the depocentre had migrated towards the NE, and a normal marine faunal assemblage had returned. A notable difference with the Aquitanian sequences was the presence of well-developed coral build-ups along the platform margin. During the early Miocene, carbonate was the dominant lithology in the majority of the shallowwater environments. Mixed carbonate –siliciclastics were only found in the most proximal locations in the southwestern part of the study area, whereas in the intrashelf basin deeper marine marls and anhydrites were deposited. Sequence 4 (early Aquitanian). The boundary at the base of this sequence (SB IV) is described in the previous section. The top sequence boundary (SB V) is picked in the basin centre at the base of the ‘Middle Anhydrite’ (e.g. Well-13 in Figs 10 & 17b). At the margin SB V is picked on top of a shallowing upward trend (e.g. Kuh-e-Asmari; Fig. 9), while at the platform top the boundary is based on a combination of sedimentological evidence and Sr isotope dating (Fig. 14). In the sandstone dominated wells of the southwestern area the boundary has been picked on thickness, in analogy with adjacent carbonate dominated wells where this surface has been dated with Sr isotopes (Fig. 17a). Since the overlying sequence 5 is mostly absent on top of the platforms, SB V is interpreted as a subaerial exposure surface (Type I sequence boundary). The lithological composition of this sequence is mixed in the southwestern part of the basin, carbonate dominated in the northeastern and southeastern part, and mixed anhydrite/marls/carbonates along the margins and in the basin centre in the NW

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(Fig. 17). The lowstand deposits of this sequence consist of a basinal anhydrite bed (the ‘Basal Anhydrite’) ranging in thickness from 3 to 10 m. Deep water basinal deposits with a pelagic fauna were found below and above these anhydrite deposits in Kuh-e-Asmari, Well-11, Well-12 and Well-13 (Fig. 17). Along the margins of the basin the anhydrites are intercalated with shallow-water carbonates (e.g. Kuh-e-Bangestan; Figs 9c & 17b). The deposition of basin-wide subaqueous anhydrites is interpreted as evidence for the isolation of the intrashelf basin, probably caused by a big eustatic sealevel drop (see below). In the southwestern part of the study area a thin layer of sandstones is observed at the base of this sequence, overlaying an exposure surface (SB IV; Fig. 17). These sands are interpreted as the (terrestrial) lowstand deposits on the platform top, which were probably reworked during early transgression. Since no siliciclastics have been observed along the margin, or in the intrashelf basin, there was probably no bypass of siliciclastics. The high resolution dating obtained with the Sr isotopes allowed the precise dating of the MFS in this sequence (Figs 13 & 14). This shows that during transgression the carbonate platforms aggraded strongly (e.g. Well-9 and Kuh-e-Bangestan in Fig. 14), and that the sandstones in the SW backstepped and accumulated in the proximal domain (Fig. 17b). Oolitic limestone beds have been recorded at several locations on the platform top in the TST (Kuh-e-Khaviz, Kuh-e-Khami, Kuh-e-Bangestan; Fig. 17b). The continuity of these oolitic beds is difficult to evaluate, but in any case this seems to indicate that this was a phase with widespread oolite development on the platform. The HST is essentially deposited in the intra-shelf basin, as regressive prograding carbonate wedges (Well-11, Well-12, Chidan and Kuhe-Asmari; Fig. 17b). The carbonate facies was characterized by the mono-specific dominance of the crustacean coprolite Favreina (Assemblage D), which testifies to periods of environmental stress such as elevated salinity of the water in the isolated basin. In platform areas in the SW and SE, sedimentation was condensed or the platform was subaerially exposed, such as is suggested in Transect 1 (Fig. 17a). The Sr-isotope dating helped to identify a clear volumetric partitioning, distinguishing between the transgressive phase of the sequence, when the platform aggraded and little sedimentation occurred in the basin, and the regressive phase of the sequence, when sedimentation shifted to the intrashelf basin, with condensed sedimentation on the platform (Fig. 14). Sequence 5 (late Aquitanian). The base of this sequence, SB V, is presented in the previous section. The top boundary, SB VI, is in most of

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the platform area amalgamated with the underlying SB V (Fig. 17a), and interpreted as a subaerial exposure surface (Type I sequence boundary). In the intrashelf basin both surfaces are clearly separated (e.g. Well-13, Fig. 17b), and also locally along the margin of the basin (e.g. in Well-11, Well-12, Kuh-e-Asmari, Chidan, Izeh; Fig. 17a). This surface marks the boundary between the Aquitanian and Burdigalian, and is well constrained by both the occurrence of the benthic index fossil Borelis melo curdica in the Burdigalian and the Sr isotope dating (Figs 13 & 14). The lithology in this sequence consists of limestones, anhydrite, marls, and only a little sand. The lowstand deposits consist of a well developed anhydrite interval (‘Middle Anhydrite’) that is present in the basin centre (Well-13, Fig. 17b). In Well-13, the anhydrite is overlying pelagic intrashelf basinal marls, and reaches a maximum thickness of 50 to 70 m, showing a vertical evolution from one or two decametre thick beds at the base, to thinner metre-scale beds at the top that grade into an argillaceous limestone (Fig. 10). Lateral facies change is observed from the massive anhydrite beds in the basin centre, to a more heterogeneous composition along the margin where interfingering with the carbonates occurs. A sharp eustatic sea-level drop is invoked to isolate the basin and create conditions for the precipitation of the subaqueous anhydrites. The gradual vertical facies evolution, suggests a stepwise sea-level rise, possibly controlled by higher order eustatic sealevel fluctuations, which gradually re-established a connection with the open ocean, and the return to normal carbonate deposition. The transgressive and highstand deposits of this sequence are restricted to the basin depression and its immediate borders (Figs 14 & 17). In the outcrops of Kuh-e-Asmari and Chidan a restricted, locally dolomitized facies is found. The rest of the platform was exposed at that time, evidence for which has been found in core material of, for example, the Ahwaz field (palaeosoils). At the end of this sequence the intra-shelf basin was filled to spill, creating a flat, partly exposed platform in the study area (Fig. 17). Sequence 6 (early Burdigalian). The lower sequence boundary (SB VI) has been presented in the previous section. The top boundary, SB VII, is picked below the first anhydrite of the overlying Gachsaran Formation. Sr isotope dating of this anhydrite has given consistent ages, suggesting that this is, indeed, an isochronous event, and can thus be picked as a timeline (Fig. 13 and Table 3). The lithological composition of this sequence consists of shallow-water limestones and basinal marls. No lowstand deposits have been identified

in this sequence. The gross depositional environment during sequence deposition consisted of a large flat platform with a seaward margin located in the NE (Fig. 17a). The sequence commenced with the deposition of shallow-water deposits in most of the area, suggesting the entire platform was, at least briefly, flooded during the early TST. In the most northeasterly sections these shallowwater deposits are abruptly overlain by pelagic marls (e.g. Katoola section; Figs 8 & 17a), while, in the same time interval in the Chidan section, coral buildups were deposited. Cross-sections show a sharp increase in sequence thickness towards the NE, and a thinning towards the south, where locally no sediment was deposited (Fig. 17a). The deepening in the north is interpreted as the result of a regional tilting that started in the early Burdigalian. The northeastern region rapidly subsided (270 m in Katoola) and was flooded by the sea, whilst condensed sedimentation took place in the SW caused by synsedimentary uplift. The steeply rising curves of the Sr-isotope measurements testify to the rapid subsidence and high rate of sediment accumulation at that time (Figs 13 & 14). In the majority of the studied outcrops and wells an anhydrite bed, several metres thick, is found at the top of this sequence, marking the base of the Gachsaran Formation. Sr isotope dating of this anhydrite bed has given consistent ages (Fig. 13), suggesting this was a virtually isochronous event. This environmental change was coincident with the closure of the NeoTethys Ocean.

Discussion Age dating In this study a systematic integration of biostratigraphic and chemostratigraphic methods has been applied to improve the age dating of the OligoMiocene in the Dezful Embayment. This age dating has been instrumental in: (1) the establishment of a basin wide, high resolution set of timelines; (2) the recalibration of the stratigraphic ranges of a number of benthic foraminifera index fossils; (3) the revision of the age assignment of a number of rock units; and (4) demonstrating the dynamics of the sedimentary system (facies partitioning and sedimentation rates). The current study clearly shows that the old age interpretation based on large foraminifera needs revision. One of the fundamental changes is that biozones previously attributed to the Aquitanian are in fact Chattian in age (Fig. 16). Also, the study clearly showed that so called ‘marker species’, especially of the Wynd (1965) zonation, are extremely facies dependant and should not be

STRATIGRAPHIC ARCHITECTURE AND RESERVOIR TYPES OF THE ASMARI FORMATION

used for dating purposes. For example, Peneroplis evolutus, marker of the previously Aquitanian ‘Austrotrilina howchini – Peneroplis evolutus Assemblage Zone’ of Wynd (1965) ranges from the Rupelian to Burdigalian in our dataset (Fig. 15). The zonation of Adams & Bourgeois (1967) on the other hand is fairly robust and the marker species more stratigraphically restricted. Herein, it is proposed that the age assignment of the ‘Archaias asmaricus – Archaias hensoni Assemblage Subzone’, previously assigned to Aquitanian by Adams & Bourgeois (1967), be changed to Chattian (Fig. 16). In the old age interpretation Chattian was not separated from the Rupelian, and there was just a general ‘Oligocene’ age assignment. The current study shows that we can in fact separate Chattian, as a part of the previously interpreted Aquitanian, and Rupelian, the interval previously stated to be ‘Oligocene’ (Figs 15 & 16). The limited dataset, with Nummulites observed at 4 localities, shows that the presence of Nummulites species indicates a Rupelian age. Cahuzac & Poignant (1997) indicate that the species could range into the Chattian, but in our dataset the top occurrence is well within the Rupelian. The widespread extinction of this marker at the end-Rupelian elsewhere supports this conclusion. Good Chattian markers are Archaias species, as they generally seem to become extinct at the Chattian/Aquitanian boundary (except for the northeastern corner of our dataset, which is possibly caused by the low density of Sr isotope samples in the Katoola section). Miogypsinoides complanatus is another valid Chattian marker. Based on the limited data in this study, and on its range in the Indo-Pacific, Spiroclypeus is potentially a useful Chattian marker in the Asmari Limestone, but this requires further analysis. The Aquitanian is more difficult to identify with marker species. A general observation was that when Miogypsina spp. and Elphidium sp. 14 occur together, an Aquitanian age is likely. Favreina asmaricus appears to be an excellent marker for the Aquitanian, however, it is only present in the northern part of the dataset. The particular environmental conditions during the Aquitanian (isolated hypersaline basins for part of the time) may play a role in the occurrence of this coprolitic facies. The Burdigalian is easily picked with the base occurrence of Borelis melo curdica (exception is in Gachsaran where strontium ages indicated older ages). Locally, the Sr isotope stratigraphy provided a high enough precision to reveal the dynamics of the sedimentary system within a depositional sequence. The shape of any particular Sr age curve reflects changes in sedimentation rates, and hence complex geometries related to depocentre shifts and volumetric partitioning can be clearly

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demonstrated. A good example was found in sequences 4 and 5, where in Well-9 a fast aggrading TST of sequence 4 was observed, whereas the HST of this sequence and almost all of sequence 5 are condensed in a plateau of the curve (Fig. 14). On the other side of the basin, in the Izeh section, a comparable condensed section was observed, illustrating that sequences therefore prograded into the basin from two sides (Fig. 14). In Burdigalian sequence 6 a regional thinning is observed from the NE to the SW (Figs 14a & 17a). Using the Sr isotope curves this can be interpreted as being caused by a regional differential subsidence causing condensation in the SW, and expansion of the stratigraphic succession in the NE, rather than erosion. The structural process is interpreted as a synsedimentary regional flexure towards the NE.

Controlling factors The relative influence of tectonics on the depositional history can be evaluated using the two regional transects. Both have been flattened on SB VI, the base of the Burdigalian, when the basin was filled, and water depth showed little variation, ranging from subaerial exposure to very shallowwater marine (Fig. 17). In both transects the Maastrichtian –Paleocene boundary forms the base of the studied succession, which allows evaluating the total amount of accommodation created from the Paleocene to the early Miocene. The following observations about the basin dynamics can be made: (1) during the Eocene differential subsidence occurred with more accommodation created in the SW where a very thick succession of shallow-water platform deposits of the Jahrum Formation accumulated; (2) during the Oligocene sedimentation took place in the under-filled basin confined by the palaeo-relief of the subaerially exposed Jahrum carbonate platform. Transect 1 shows, however, a regional variation in the accommodation space within this basin: along the basin margins in the southwestern part, and to a lesser extent in the southeastern part, a higher subsidence is observed along the Jahrum Platform, while a lower subsidence area is observed towards the middle of the basin (Well-10 and Well-11 in Fig. 17a). Sediment accumulation in this phase varied from a maximum thickness of approximately 350 m along the basin margins to a minimum thickness in the basin of approximately 100 m; (3) during the Aquitanian sedimentation was mostly confined to the basin, locally reaching a thickness of up to 250 m (Well-13; Figs 10 & 17b), while only little (and locally no) sediment accumulated on the platforms. No evidence for Aquitanian synsedimentary tectonics has been observed in the study area; (4) during the Burdigalian regional tectonic tilting

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towards the NE controlled the sedimentation in the study area, with a maximum sediment accumulation of 270 m in the NE, and locally condensed sedimentation or exposure in the SW. From these geometrical considerations it can be concluded that the maximum depth of the Oligocene/Miocene basin in the study area was in the order of 250 m or less. Global climatic cooling and Antarctic glaciation influenced sedimentation worldwide during the middle Eocene to early Miocene. This has been documented in marine and terrestrial faunal and floral changes, seismic interpretations and oxygen and carbon isotope analyses (e.g. Prothero 1994; Miller et al. 1991, 2005; Abreu & Haddad 1998; Zachos et al. 1999, 2001; Clarke & Jenkyns 1999). Three significant climatic cooling and glaciation events are particularly relevant for this study: at the Eocene –Oligocene boundary, in the middle Oligocene (Rupelian/Chattian boundary), and at the Oligocene –Miocene boundary. In addition, within the Aquitanian there is evidence for yet another cooling event (Miller et al. 2005). The sequence subdivision proposed here and the associated lithological changes appear to coincide with these global variations in climate and sea-level. Firstly, there is good evidence for exposure of the Jahrum platform during the late Eocene and early part of the Oligocene with the onlap configuration of the Rupelian carbonates. The appearance of Rupelian shallow-water deposits in the northeastern part of the basin at this time may also be (partially) related to this sea-level drop. Secondly, the Rupelian-Chattian boundary is characterized in the Dezful Embayment by the sudden, massive influx of siliciclastics from the hinterland in the SW, while the sequence boundary (SB II) is an exposure surface on the surrounding platforms. Thirdly, the anhydrite deposits in the centre of the isolated intrashelf basin (lowstands of sequences 5 and 6), occurring around the Oligocene –Miocene boundary and in the middle of the Aquitanian, coincide again with reported sea-level drops. A difference between the Oligocene and Aquitanian lowstands is the deposition of the anhydrites in the basin during the Aquitanian. This may be explained by the combination of a well established carbonate platform barrier, which became exposed during sea-level fall, and the generally warmer (arid) climatic conditions during the Aquitanian (e.g. Prothero 1994; and references therein), causing the precipitation of subaqueous anhydrites from the oversaturated seawater through evaporation. In more detail, there appear to be some differences in the amplitude of the sea-level changes observed in the Dezful embayment and those reported for the global sea-level curves (cf. Miller et al. 1991, 2005; Abreu & Haddad 1998). Sea-level fluctuations in the Dezful Embayment were

estimated using the regional geometrical reconstructions, the sedimentological facies interpretations and the high-resolution timelines. The latest Eocene sea-level drop (SB-I) was the most important one, with an estimated fall of at least 100 m. Subsequent falls were smaller, and probably not bigger than 30 –40 m. The magnitude of sea-level rise varied also, with the biggest rise occurring during the early Chattian, as is clearly illustrated by the clinoforms in Tang-e-Gurguda, that downlap on the underlying sequence with shallower water facies. Sea-level dropped down the platform top again twice in the Aquitanian (the periods of the anhydrite deposition), and sedimentation was mostly limited within the intrashelf basin with very little sediment accumulating on the shelves, suggesting a relatively small eustatic sea-level rise (20– 30 m) at that time.

Comparison with other sequence stratigraphic interpretations Previous sequence stratigraphic analyses of the Oligo-Miocene deposits in SW Iran include the regional work by Sharland et al. (2001, 2004), and the more local studies by Ehrenberg et al. (2007) and Vaziri-Moghaddam et al. (2006). In particular the first two are of relevance for this study. Ehrenberg et al. (2007), in their study of the Asmari Formation in the southwestern part of the Dezful Embayment, defined 7 surfaces that were dated and correlated between the Ahwaz, Marun and Bibi Hakimeh Oilfields and a set of outcrops in the Kuh-e-Khaviz Anticline. Six of the defined surfaces were considered to be major sequence boundaries of which five can be matched with the sequence boundaries defining the base of each sequence described herein (Fig. 19). In increasing age they are: Bu20 Sequence Boundary of Ehrenberg et al. (2007) equates with SB VI; Aq20/Bu10 with SB V; Aq10 with SB IV; Ch30 with SB III and Ru30/Ch10 with SB II. Surface I in the current study had no counterpart in Ehrenberg et al. (2007) since neither the cores nor outcrop locations penetrated or exposed such old strata. Of the remaining 2 surfaces, the Ch20 of Ehrenberg et al. (2007) has no counterpart here and the intra-Aq10 surface that they mapped was dated some 500 000 years younger (23.0 Ma) than the Aq10 surface (23.5 Ma) and is probably a prominent fourth order depositional sequence that could be mapped over their study area. One of the interesting features in utilizing Strontium isotopes to date the regional sequence boundaries is that certain surfaces, particularly those at the base of Sequence 4 (SB IV) and 5 (SB V) show a pattern of decreasing age toward the basin centre (Fig. 14). This feature

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Fig. 19. Comparitive scheme for sequences. Comparison of sequences and age assignments for European Basins (Hardenbol et al. 1998), the Arabian plate (Sharland et al. 2001, 2004), and the Dezful Embayment (Ehrenberg et al. 2007, and this study). See text for discussion.

was also noted by Ehrenberg et al. (2007) in their corresponding Aq10 and Aq20/Bu10 sequence boundaries. Both of these depositional sequences are believed to have been strongly progradational and in such a situation the age of strata exposed below the surface will young towards the basin centre as the component high-order depositional packages possess offlap and eventually down-step geometries. Sharland et al. (2001, 2004) in their study of the Arabian Plate defined a set of regional Maximum Flooding Surfaces (MFS) that they use to subdivide the Arabian Plate sedimentary succession into a series of isochronous packages. These Genetic Stratigraphic Sequences (sensu Galloway 1989) were placed into a series of Tectonostratigraphic Megasquences (TMS) that subdivided the geological history of the Arabian Plate from Precambrian

to the Present (Sharland et al. 2001 & 2004). The Oligocene –Miocene aged Asmari, Pabdeh and Gachsaran Formations sit within the latest of the TMS, the AP11, which is defined as the package of sediments lying between the unconformity marking both the onset of Red Sea Rifting and the first continental collision between Arabia and Eurasia and the present day topographic surface (Sharland et al. 2001). Their Pg30 flooding surface has Nummulites fichteli as its index fossil and is described as occurring close to the base of the Asmari Formation in Iran (Sharland et al. 2001). Lepidocyclinids are not associated with this flood and this surface would therefore coincide with the MFS of Sequence 1 in this study (Fig. 19). This MFS was dated to 33 Ma by Sharland et al. (2001) on the time-scale of Gradstein & Ogg (1996), which corresponds to approximately 32.8 Ma on

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the Gradstein et al. (2004) time-scale and it is in keeping to what is found in this study. The Pg40 flooding surface was added by Sharland et al. (2004) and is dated to 29 Ma on the time-scale of Gradstein & Ogg (1996), which corresponds to a date of 29.1 Ma on the Gradstein et al. (2004) scale. It is characterized by the presence of lepidocyclinids and would therefore correspond to the MFS of Sequence 2 in this study, though Sr dating in the Asmari would indicate a somewhat younger age. Pg50 MFS, formerly the Ng10 MFS of Sharland et al. (2001), was originally assigned to a shale unit in the Ab Teymur well but was subsequently moved to the overlying carbonate interval (Sharland et al. 2004). The diagnostic fossil for the Pg50 is Miogypsinoides complanatus and the MFS has been assigned an age of 24.5 Ma (Sharland et al. 2004) on the Gradstein & Ogg (1996) time-scale, which would equate to approximately 23.5 Ma on Gradstein et al. (2004). The Pg50 MFS therefore corresponds to the Sequence 3 MFS in this study. Based on the improved age control and stratigraphic understanding a revision of the Ng10 MFS of Sharland et al. (2001, 2004) is proposed here. With the absence of Sr isotope dating in Iraq, the most reliable correlation guidelines are the anhydrite deposits. In the Iraq part of the basin, also two anhydrite levels are found: the informally known ‘Basal Anhydrite’, which lies below the Euphrates and Serikagni Formations dated as Aquitanian by van Bellen et al. (1959) and Starkie (1994), and the anhydritic Dhiban Formation, which is overlain by the Jeribe Formation, being consigned to the top of the early Miocene and of Burdigalian age (van Bellen et al. 1959; Sharland et al. 2004). In our view, the flood at the base of the Euphrates and Serikagni Formations correlates probably with Sequence 4, since these Iraqi formations overlie the Basal Anhydrite, and the Ng10 MFS would therefore sit within this sequence. It would then represent a base Aquitanian regional flood and not a Burdigalian event as is suggested by Sharland et al. (2004), and is therefore older than what they quote; c. 23 Ma as opposed to c. 20 Ma. The Ng20 has a distinctive fauna characterized by the presence of Borelis melo curdica and has been dated to 18.5 Ma by Sharland et al. (2004) and 17.5 Ma by Simmons et al. (2007). In Iran the Ng20 is placed by Sharland et al. (2004) in the Upper Asmari carbonates as defined by Slinger & Crichton (1959) and Motiei (1993) and in the planktonic limestones described within the Burdigalian succession of Seyrafian & Hamedani (1998). In Iraq it occurs at the base of the Jeribe Formation, above the Dhiban anhydrite (Sharland et al. 2004). The Ng20 would thus most likely correspond to the flood at the base of our Sequence 6, although it

is dated somewhat older to between 18.7 and 19.0 Ma on the Gradstein et al. (2004) time-scale, which equates to between 18.8 to 19.1 Ma on the Berggren et al. (1995) scale. In conclusion, this paper is in general agreement with the sequence subdivision proposed by Ehrenberg et al. (2007). Compared to the more regional work of Sharland et al. (2001, 2004), this paper provides good evidence for one additional, regionally ‘robust’ sequence in the Aquitanian (sequence 5, a possible Ng15). This is an important sequence, considering the regional anhydrite deposition that took place during its lowstand. In addition, a revision of the age assignment of Ng10 is proposed to the early Aquitanian (Fig. 19).

Regional implications The Asmari Formation in the Dezful region forms just one small part of a very extensive OligoMiocene carbonate province that was deposited in a foreland basin that can be traced from the Fars province in Southeast Iran to Syria. Oligo-Miocene carbonates are seeded on both sides of this basin on former Eocene-aged platform carbonates or occasionally on mixed carbonate-siliciclastic successions, such as in the northeastern side of the basin in Lurestan where the Asmari may sit directly on either Eocene aged carbonates of the Shahbazan Formation or possibly directly on Paleocene to ?Eocene aged red-beds of the Kashkan Formation (Homke et al. 2004). Along the entire southern side of the basin from Syria to Iran, the Asmari and its lateral equivalents in Iraq are seeded onto the former Eocene platforms (Damman Formation in Iraq and the Jahrum in Iran). The Asmari is actually equivalent to 11 different formations in Iraq; in broad terms Oligocene-aged Asmari strata are correlated with the 9 formations that comprise the Kirkuk Group, whereas Aquitanian-aged strata are equated to the Euphrates Formation and Burdigalian-aged strata with the Jeribe Formation (cf. van Bellen et al. 1959). However, it must be noted that these correlations are by no-means as straightforward as might first appear and in the absence of detailed published biostratigraphic sections, and no Sr-isotope dating, identification of the main depositional sequences in the Iraqi Oligo-Miocene stratigraphy remains tentative (see above). The following, robust, sequence stratigraphic model can be proposed for the region covering the Dezful Embayment, Lurestan and eastern Iraq: Rupelian/Chattian. The latest Eocene major sealevel drop set the scene for the Oligocene deposition. The first Oligocene sequence (Sequence 1 of Rupelian age) was deposited in an onlap position against the exposed Eocene platforms. The Rupelian

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platforms were characterized by the presence of Nummulites, and no siliciclastics were deposited in the Dezful Embayment. Sequence 2 (early Chattian), marks the influx of large amounts of siliciclastics into the Dezful Embayment from the SW. These sands either (a) by-passed the exposed platform and were deposited as lowstand shallow marine systems sitting out-board of the carbonate platforms; or (b) were trapped during high-order rises in sea-level on the platform top as transgressive parasequences. In the northeastern margin pure carbonate sedimentation occurred. Progradation into the basin occurred from both sides. Sequence 3 (late Chattian) continued in a similar style to the previous sequence, with a mixed carbonate –siliciclastic system in the SW and a carbonate system in the NE. An intrashelf basin gradually became isolated from the open ocean, and was filled by progradation of the basin margins. Aquitanian. Sequence 4 (early Aquitanian) is characterized by the first deposition of subaqueous evaporites (the ‘Basal Anhydrite’) during the sequence lowstand. In Lurestan this anhydrite unit increases significantly in thickness and it is possible that halite might have accumulated (lower Kalhur Member). In Iraq this unit is also informally known as the ‘Basal Anhydrite’. Indeed, this unit of anhydrite can be traced from Khuzestan in Iran to Syria, suggesting the presence of one large basin that became temporarily disconnected from the open ocean. During the following sea-level rise, the widespread occurrence of the coprolite Favreina suggests that crustaceans were thriving in this isolated basin with a stressed environment caused by a relatively high salinity. Sequence 5 (late Aquitanian) is characterized by a second phase of anhydrite deposition (the ‘Middle’ Anhydrite). In Lurestan the Middle Anhydrite reaches a thickness of c. 150 m and is called the Kalhur Member and it often contains at least one thick halite unit. In Iraq this interval correlates with the Dhiban Formation, which also contains halite interbeds in the basin centre (Al-Juboury et al. 2007). During this sequence the platforms were mostly exposed, and did not record much sediment. Channel incision probably occurred at this time in the proximal settings in the SW. By the end of this sequence the basin topography in Khuzestan was mostly filled in, and the coastline shifted towards the NE (Izeh area). Burdigalian. In sequence 6, thick deposits of shallow-water carbonates followed by marine marls were deposited in the Izeh zone, while condensed sedimentation, non-deposition and incision took place in the southwestern areas. This sequence illustrates most clearly evidence for a strong tectonic influence on the sedimentation pattern in the Dezful Embayment (regional tilting). The impact of this event in Iraq needs further study.

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Reservoir implications Four reference models of Asmari reservoirs are proposed based on the variations in the stratigraphic architecture (geometries and facies variability) and the lithological composition observed in the Dezful Embayment (Fig. 20). These are: Type 1, a siliciclastic dominated system; Type 2, a mixed carbonate –siliciclastic system; Type 3, a mixed carbonate –anhydrite system; and Type 4, a carbonate dominated system. These reference types have specific geographical distributions that are strongly controlled by the palaeogeography of the basin: the siliciclastic and mixed siliciclastic-carbonate systems are located in the western part of the area, the mixed carbonate–anhydrite system in the central part of the intrashelf basin, and the carbonate dominated in the eastern and southern part of the basin (Fig. 20). In Table 4 the main stratigraphic characteristics of these four reservoir reference models are listed. It is clear that the stratigraphic architecture and lithological composition are only two of the factors that control reservoir characteristics, and that subsequent diagenetic processes and structural deformation have further altered the reservoir characteristics (e.g. McQuillan 1974; Wennberg et al. 2006, 2007). A few general statements, however, can be made about the (potential) influence of a diagenetic overprint on these four stratigraphic reference models. In Type 1 (siliciclastic dominated) relatively little influence of diagenesis has been observed, due to the dominance of the siliciclastic fraction and virtual absence of carbonate and clayey interbeds. Any diagenetic overprint will be strongly controlled by the variation of permeability in sandstone bodies. In Type 2 (mixed carbonate-siliciclastic) the mixed lithology could lead to a number of diagenetic overprints. Both early and burial dolomitization could be expected in the carbonate facies, whilst sandstone facies may be carbonate cemented. There is also a potential for burial diagenetic overprint in both the carbonate and sandstone facies. In Type 3 (mixed carbonate-anhydrite) the presence of anhydrite layers may lead to remobilization of sulphate during burial and thus to anhydrite plugging of the surrounding carbonates. In Type 4 (carbonate-dominated) early dolomitization processes can be expected, in particular in the muddy carbonate facies, which may have been subsequently overprinted by burial diagenesis. The combination of stratigraphic architecture, lithological variation and diagenetic overprint control the rock mechanic characteristics that influenced fracture patterns. The relationship between the different diagenetic phases and the fracturing styles is presented by Daniel et al. (2006, 2008) and Ahmadhadi et al. (2008).

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Fig. 20. Stratigraphic reference models for the Asmari reservoir types. This study has led to the proposition of four types of Asmari reservoirs (see Table 4). Their lithological, stratigraphic and geographical characteristics are summarized in this figure. Whereas their basic distinction is a lithological classification, the lateral changes and geometries that characterize them can be predicted using the sequence stratigraphic model presented in this paper.

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Table 4. Stratigraphic reference models for Asmari reservoirs Stratigraphic reference model 1 (siliciclastic dominated) – Lithologies: sandstone dominated, with clayey interbeds, locally thin carbonate layers (towards the top). – Geometries: decametre thick accumulations of flat bedded continuous sands, internally metre and decimetre scale heterogeneities, local incisions, reactivations possible. – Distribution: in the western, more proximal part of the study area, including sites located on top of the Eocene Jahrum platform. – Age: mostly sequences 3 and 4 (late Chattian, early Aquitanian), locally also sequences 2 and 6 (early Chattian, Burdigalian). – Diagenetic overprint and fracturing: sandstone setting, diagenesis is not well known. – Fields/structures: e.g. Azadegan, Dorquain, Jufeyr, Korramshahr, Kushk, Susangerd. Stratigraphic reference model 2 (mixed carbonate-siliciclastic) – Lithologies: mix of shallow-water carbonate deposits and siliciclastics; siliciclastics are best developed in the west, and gradually wedge out towards the east and the north. – Geometries: sandstones are deposited as sheets on top of the carbonate platforms, and in a more heterogenous manner, as lobes or bars along the margins of the intrashelf basin (cleaning up trends in the GR logs), or as turbidites in the basin center. The carbonates that are deposited in the platform margin, can be massive and grainy and deposited in prograding clinoforms. Platform tops are characterized by metre to decimetre scale bedding, with a large lateral continuity. – Distribution: this model occurs along the western margin of the intrashelf basin, locally also on top of the Eocene Jahrum. – Age: sequences 2 to 6 (Rupelian to Burdigalian) – Diagenetic overprint and fracturing: the mix of sandstones and carbonates may control both the style of diagenesis and fracturing in these rocks – Fields/ structures: e.g. Agha Jari, Ahwaz, Kilur Karim, Kupal, Marun, Rag-e-Sefid. Stratigraphic reference model 3 (mixed carbonate-anhydrite) – Lithologies: mix of carbonates and anhydrites. The anhydrites are confined to the intrashelf basin centre, where they are encased in basinal marls, while towards the margins they interfinger with the shallow-water carbonates. – Geometries: anhydrites occur in two horizons (lowstands of sequence 4 and 5) and wedge out laterally into the shallow-water carbonates. At the time of deposition of the anhydrites the basin was temporarely isolated from the open sea (NeoTethys). – Distribution: this mixed system occurs only in the central part of the intrashelf basin (during sequences 4 and 5). – Age: the base of sequences 4 and 5 (Aquitanian). – Diagenetic overprint and fracturing: the mix of carbonates and anhydrites creates a geochemically reactive environment. The lithological heterogeneity may affect both the large scale deformation (detachment levels in the anhydrites), as well as the fracture pattern. – Fields/structures: e.g. Haft Kel, Kuh-e-Asmari, Kuh-e-Bangestan, Lali, Papileh, Zeloi. Stratigraphic reference model 4 (carbonate dominated) – Lithologies: in this model the sediment consists for more than 90% of carbonate, only very locally sandstone facies have been observed (e.g. offshore sands in Kuh-e-Khaviz, sand layer in Kuh-e-Khami, sand layer in Katoola), – Geometries: Depositional geometries in this system are variable and vary from large scale clinoforms (with 90 metres of relief), to decametre scale coral buildups (laterally kilometre scale), and well bedded (metre to decimetre scale) platform top deposits. – Distribution: along the northeastern margin of the basin. – Age: sequences 1 to 6 (Rupelian to Burdigalian) – Diagenetic overprint and fracturing: in the carbonate environment both primary and burial diagenetic overprint are present; fracturing is a known characteristic of the carbonate fields. – Fields/structures: Bibi Hakimeh, Gachsaran, Katoola, Kuh-e-Mish (Tang-e-Gurguda), Mansurabad, Pazanan.

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Conclusions The following conclusions are drawn:

Improved time framework Sr isotope stratigraphy has been systematically applied to date the outcrop sections and the cored wells presented in this study. This dataset has recalibrated the ranges of a number of benthic organisms that have been used in the past to date the Oligo-Miocene shallow-water successions. This has led to a revision of the biostratigraphic zonation scheme for the Oligo-Miocene in the Dezful Embayment, with notably a change in the position of the Chattian–Aquitanian boundary, and a refinement of the zonation in the Oligocene, with the distinction between Rupelian and Chattian. The improved time control allowed the identification of timelines across different depositional environments, and the comparison of the local sedimentation pattern to worldwide events such as glacio-eustatic sea-level fluctuations and climate changes.

Basinwide sequence stratigraphic model Due to the improved time control a robust framework of six depositional sequences has been constructed, including sequence boundaries and maximum flooding surfaces. These sequences have been cross correlated from platform to basin through 50 control points in the study area (300  350 km), which allowed unravelling the complex 3D geometrical and lithological organization of this Oligo-Miocene sedimentary system. These six sequences represent a total period of 15.4 Ma, and vary in duration from 1.2 to 5.2 Ma, which classifies them as third order. Three Oligocene sequences are distinguished: Sequence 1, of latest Eocene to Rupelian age, was deposited in an onlap position against the exposed Eocene Jahrum platform in the SW. This shallowwater carbonate platform prograded in a northeastly direction. Sequence 2, of early Chattian age, was also partly deposited in onlap position against the Jahrum palaeo-relief, and was characterized by the onset of a major influx of siliciclastics from the SW. During sequence 3, of late Chattian age, mixed siliciclastic sedimentation continued in the SW, and carbonate-dominated sedimentation along the northeastern margin. This sequence was terminated with the major glacio-eustatic sea-level drop that occurred at the Oligocene–Miocene boundary. Three Miocene sequences are distinguished: Sequences 4 and 5, of Aquitanian age (early Miocene), both started with lowstand deposits of subaqueous anhydrites (‘Basal Anhydrite’ and ‘Middle Anhydrite’) that were deposited when the basin became isolated from the NeoTethys ocean. Limited accommodation on top of the carbonate

platforms during transgression and highstand forced sedimentation into the intrashelf basin, where a confined environment existed, as suggested by the dominance of the Favreina-rich facies. Sediment supply was carbonate-dominated. By the end of sequence 5 the basin was filled to spill, creating a flat, partly exposed platform in the study area. Sequence 6, of Burdigalian age (middle Miocene), represents the last phase of the Asmari deposition. Regional tectonic tilt caused exposure and nondeposition in the SW, and marine transgression and increase in accommodation in the NE. The sequence was terminated by a regional shallow-water evaporitic environment (Gachsaran Formation), which marked the closure of the NeoTethys at that time.

Four stratigraphic reference models for Asmari reservoirs Based on lithological composition and stratigraphic organization four stratigraphic reference types are defined which each have a different palaeogeographical distribution: Type 1, sandstone dominated system; Type 2, mixed carbonate–siliciclastic system; Type 3, mixed carbonate–anhydrite system; Type 4, carbonate dominated system. This lithology based classification gains in predictive power through the association with the sequence stratigraphic model that predicts facies trends in time and space. For example, the carbonate-dominated and mixed carbonate –siliciclastic reference types are diachronous in a basinward direction due to the progradational nature of the sedimentary system. In addition, the mixed carbonate-anhydrite system is only present in the Aquitanian, when the basin became isolated. Superposed on these stratigraphic reference models are the diagenetic processes and structural deformation (faulting, folding and fracturing) that will eventually determine the specific reservoir characteristics. The R&D Directorate of the NIOC initiated in 2004 an industry financed research project to study EOR aspects of the Asmari Reservoirs. The re-evaluation of the stratigraphic architecture of the Asmari Formation in the Dezful Embayment was one of the objectives of this project, and results presented here were achieved between 2004 and 2006. Support of the management of the Exploration Directorate, the IOOC, the NISOC, the RIPI are acknowledged for this project by providing access to data and participation of their staff members. We are grateful to the management of the EOR project, and in particular Mr S. Hamzai (NIOC R&D), Mr E. Garborg (Statoil) and Mr G. Drullion (Total) for their motivation, enthusiasm and perseverance, to make this study happen. The EOR project management and the financing companies Petronas, Shell, Total, Norsk Hydro/Statoil are acknowledged for permission to publish this paper. We wish to thank F. Gaumet and A. Ghobeshavi for support in the field, and C. Griffith, S. Lopez and D. Granjeon for discussions on the stratigraphic modelling of this system.

STRATIGRAPHIC ARCHITECTURE AND RESERVOIR TYPES OF THE ASMARI FORMATION We wish to thank A. Hosseini and F. Nasrabadi in Tehran, and N. Doizelet at the IFP for drafting support. Reviewers, K. D. Gerdes and A. Horbury, helped to substantially improve this manuscript.

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Architecture and facies differentiation within a Middle Miocene carbonate platform, Ermenek, Mut Basin, southern Turkey X. JANSON1,2,3*, F. S. P. VAN BUCHEM1,4, G. DROMART5, H. T. EICHENSEER2, X. DELLAMONICA2, R. BOICHARD2, F. BONNAFFE3,6 & G. EBERLI6 1

Institut Franc¸ais du Pe´trole, Geology-Geochemistry division, 1-4 Avenue de Bois Pre´au 92852 Rueil-Malmaison, France 2

Total, CSTJF-Avenue Larribau, 64018 Pau cedex, France

3

Present address: Bureau of Economic Geology, Jackson School of Geosciences, The University of Texas at Austin TX, USA 4

Present address: Maersk Oil Qatar AS,

5

Ecole Normale Supe´rieure de Lyon, France

6

Department of Marine Geology and Geophysics, University of Miami, FL, USA *Corresponding author (e-mail: [email protected]) Abstract: In south-central Turkey, a carbonate platform system of early middle Miocene age is exposed in three-dimensional outcrops displaying a rich variety of carbonate facies associated with exceptionally well-preserved depositional geometries. This paper presents a detailed reconstruction of the geometries and facies organization across the prograding margin of one intraplatform carbonate bank that grew during the Langhian on the Ermenek Platform. The total thickness of the margin is approximately 250 m, and it has prograded over a distance of 1.2 km. The geometrical pattern shows an alternation between sigmoid, sigmoid-oblique, and oblique accretionary units at different scales. Based on the facies distribution and the geometrical framework two large-scale depositional sequences and eight medium-scale depositional sequences were defined. The general evolution from a low-angle shelf geometry to a prograding flat-topped platform was associated with an evolution from oligophotic-dominated carbonate producers, such as large benthic foraminifera, molluscs, echinoderms, red algae and bryozoans at the base, to mesophotic and euphotic carbonate producer organisms, such as corals, red algae and porcellaneous small benthic foraminifera at the top. The eight medium cycles were defined primarily using the depositional geometries, since facies changes were observed only locally within these cycles. Several mechanisms influenced the stratigraphic architecture of this margin: (1) eustatic sea-level controlled the overall transgressive-regressive Langhian sequence, and two superposed large-scale sequences. Medium cycles were probably also influenced by higher frequency sea-level fluctuations; (b) climate change probably influenced the overall evolution of the faunal assemblage; and (c) antecedent topography determined the overall architecture of a shelf bordering a deeper basin.

Carbonate platform margins are one of the most sensitive depositional environments, recording the relative influence of both internal and external controlling factors on carbonate sedimentation. This makes them ideal for the study of dynamics of carbonate depositional systems, and in particular of the interplay between sediment production and accumulation. For this reason, outcrops with seismicscale geometries of carbonate margins have attracted much attention in the literature. Amongst these, Miocene carbonate exposures in the Mediterranean area are well known for their high quality and display of a large variety of stratigraphic patterns

(Franseen et al. 1993; Martin et al. 1996; Bassant 1999; Pomar 2001; Bassant et al. 2005; Warrlich et al. 2005; Janson et al. 2007). Most of these studies deal with the Western Mediterranean, and describe the combined effect of relative sea-level change and climate on the ecology and architecture of reefal carbonate deposits of Aquitanian, Burdigalian, Tortonian and Messinian age. More recently, studies have focused on the Eastern Mediterranean (Kelling et al. 2005), where in southern Turkey, there are several small Neogene and Miocene basins with well preserved carbonate platform systems (Antalya, Alanya, Mut and Adana basins; Fig. 1a).

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 265–290. DOI: 10.1144/SP329.11 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. (a) Location and geological map of the Mut Basin and the Ermenek area of South Central Turkey [simplified from Robertson (1998)]. (b) Simplified regional geological map around the Ermenek town. This map is a compilation of existing map from Demirel (1992) and Gedik et al. (1979) modified and completed by field mapping, aerial photo and satellite image interpretation.

MIOCENE PLATFORM MARGIN FROM SOUTHERN TURKEY

The study area of this paper is located in the Mut basin, which has seismic-scale, nearly undeformed exposures of early to middle Miocene carbonate reef systems (Burdigalian to Serravallian) (Bassant et al. 2005). In the western part of this basin, in the Ermenek area (Fig. 1a), 3D outcrops display a barrier reef, multiple, small pre-Miocene basement-attached platforms with small coral buildups, lagoons and a littoral zone (Figs 1b & 2). More information about this system can be found in Janson (1997, 2002) and Janson et al. (2007) where the seismic expression of this system is presented. A detailed analysis of two accretionary units in this margin can be found in Quiquerez & Dromart (2006). This paper focuses particularly on a prograding Langhian margin of one of the several small attached carbonate platform-banks that prograded into deeper intra-platform depression. This intraplatform prograding bank-margin is 250 m thick, and prograded over 1.2 km (Fig. 2b). The relationship between the well exposed depositional geometries and the facies and faunal distribution are investigated in detail to assess the relative influence on the stratigraphic architecture of the different orders of sea-level changes and the changes in carbonate productivity. Finally, the results of this study are compared to global climate and sea-level changes, which are well documented for the Langhian (Zachos et al. 2001).

Materials and methods A regional map of the eastern Mut Basin was built using a pre-existing map (Gedik et al. 1979; Demirel & Ko¨ksoy 1992; Demirel 1993), satellite photo interpretation, and field-mapping (Janson 2002; Janson et al. 2007) (Fig. 1b). A more detailed map, covering approximately 140 km2 has been established north of Ermenek town (Fig. 2a), where Burdigalian, Langhian and Serravallian lithostratigraphic units were mapped, as well as the progradation directions of the platform margins (Fig. 2a, b). Five sedimentological sections (of approximately 250 m each) have been measured along the margin transect (Fig. 2). An additional section measured by P. Bassant has also been used. To reconstruct the geometries in detail, observed surfaces on the outcrops have been recorded onto numerous photo panels and later corrected for photo distortion to construct an accurate crosssection. The biostratigraphic dating was carried out on nannofossils by C. Muller (IFP). Detailed results are reported in Bassant (1999), Bassant et al. (2005) and Janson (2002). A total of 265 thin sections have been analysed by X. Dellamonica with special attention to the foraminifera content.

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The stratigraphic terminology in this study is based on the cyclic variation of the accommodation rate versus sediment flux ratio (A0 /S0 ). The evolution of this ratio determines the depositional geometries of a margin (Van Wagoner et al. 1988; Schlager 1993). The evolution of this ratio will be displayed and symbolized by triangles. A triangle that points up indicates an increase of the A/S ratio, whereas a triangle that points down indicates a decrease of this ratio.

Regional setting, bio- and lithostratigraphy of the Ermenek platform Regional setting The Miocene was a time of extensive carbonate development around the Mediterranean. Based on the faunal assemblages and stratigraphic architecture, three phases of carbonate system development are distinguished: early Miocene (Aquitanian), early to middle Miocene (late Burdigalian and Langhian) and late Miocene (Tortonian and Messinian) (see Esteban (1996) for review). Only the two youngest carbonate phases seem to be present in the Turkish Miocene basins (Kelling et al. 2005). Here we will first introduce the study area, and then focus on the early to middle Miocene (late Burdigalian to Langhian). The Mut basin is one of several small Cenozoic basins in the southern part of the Turkey (Fig. 1a) and has recently been studied by a number of workers (Sezer 1970; Bizon et al. 1974; Koycygit 1977; Gedik et al. 1979; Korkmaz & Gedik 1990; Demirel & Ko¨ksoy 1992; Demirel 1993; Janson 1997, 2002; Broucke 1998; Bassant 1999; Le Bec 1999; Safak et al. 2005; Janson et al. 2007). Two studies are of particular relevance to this work: Safak et al. (2005), who presented a regional study of the tectonic and stratigraphy history of the Mut basin, and Bassant et al. (2005), who focused on the high resolution sequence stratigraphy of the Burdigalian and early Langhian interval (NN4 nannofossil zone). The Mut basin is located north of the Cyprus trench in the southern central part of the Taurus range, where Neogene sediments unconformably overlie deformed Mesozoic and Cenozoic rocks (Gedik et al. 1979; Williams et al. 1995). The preMiocene substratum is composed of deformed Jurassic –Cretaceous limestone of the C¸ambasitepe Formation and Cretaceous ophiolitic me´lange (Gedik et al. 1979) (Fig. 1a, b). The origin of the extension mechanism that created the intramountain basin is still currently debated. It may be related to flexural loading and faulting due to the emplacement of the Besehir-Horan nappes to the north (Williams et al. 1995). Alternatively,

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Fig. 2. (a) Geological map of the study area indicating lithofacies for main stratigraphic intervals. The arrows indicate the dip directions of depositional surfaces. For the Langhian interval, these measurements have been taken on the clinothem slope and therefore indicate progradation direction. The dashed line corresponds to a portion of the regional cross-section shown in Figure 2b. The dashed ellipse shows the inferred location of substratum topographic highs that served as nucleation points from which the carbonate platforms started to prograde out. The black polygon indicates location of Figures 5, 6, and 7. (b) Schematic cross-section showing the stratigraphic organization of the Ermenek Platform after Janson (2002). First marine deposits constitute a flat subtropical platform of late Burdigalian age. During the Langhian the shallow-water carbonate deposits consisted of an aggrading barrier that developed at the platform edge and several platform-banks that prograded into the intra-platform depression deposits. During the Serravallian, the carbonate platform consisted of an aggrading barrier and several small bioclastic shoals that developed on a flat muddy platform. The black box indicates the location of the studied margin display in Figures 5, 6 and 7.

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Fig. 2. (Continued).

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extension may have been driven by a ‘roll-back’ mechanism associated with the Cyprus subduction (see Robertson (1998) for an extended discussion on the complex tectonic setting of the region). These intra-mountain depressions were then filled during the late Oligocene to early Miocene by lacustrine and fluvial sedimentation (Demirel & Ko¨ksoy 1992; Demirel 1993). A diachronous marine transgression flooded the pre-Neogene substrata in the Mut basin and carbonate deposition began in the early Miocene in the southern part of the Mut basin, reaching the northern part in the late Burdigalian (Bassant 1999; Bassant et al. 2005). Between the late Burdigalian and the late Serravallian, carbonate sediments accumulated on the Ermenek platform. These shallow-water carbonate sediments were deposited in a barrier reef–lagoon system (Janson 1997, 2002; Janson et al. 2007). At the end of the middle Miocene, during the Tortonian, uplift began to outpace the subsidence, and this neotectonic regime of uplift has prevailed ever since, bringing Neogene deposits of the Ermenek Platform to a present-day elevation between 750 – 1750 m above sea-level (Bassant 1999; Bassant et al. 2005).

samples from the nannofossil zone NN6, of middle Serravallian age (Bassant 1999; Bassant et al. 2005; Harzhauser & Piller 2007; Piller et al. 2007).

Stratigraphic architecture of the Ermenek platform The general stratigraphic architecture of the study area is shown in the map and regional cross-section of Figure 2. The Ermenek area is located in the westernmost part of the main Mut basin (Fig. 1a). The studied area is a 20 km wide platform system that is composed of a reefal barrier in the SW, a composite platform area comprising shallow-water platforms attached to pre-Miocene substratum highs separated by intra-platform depressions, and a shore zone where the entire system onlaps onto the pre-Neogene substratum in the north (Figs 2 & 3). This carbonate-dominated system developed during late Burdigalian to Serravallian time and can be divided into three distinct intervals of carbonate platform growth based on depositional geometries, rock texture skeletal components, and nannofossil dating (see above).

Biostratigraphy and lithostratigraphy

Late Burdigalian to early Langhian: low-angle open shelf

The Miocene section in the study area starts with lacustrine deposits of Aquitanian to early Burdigalian age (dates based on foraminifera; Demirel & Ko¨ksoy 1992; Demirel 1993). These deposits make up the Yenimahalle Formation (Gedik et al. 1979; Demirel & Ko¨ksoy 1992; Demirel 1993). The subsequent early and middle Miocene marine deposits consist of the shallow-water reefal limestone Mut Formation of Gedik et al. (1979) that is dominated by red algae, benthic foraminifera, bivalves, echinoderms and corals (Fig. 2). Basinward, this limestone grades into a marly limestone with abundant planktonic foraminifera, debris of large benthic foraminifera, bivalves, echinoderms and silt-sized quartz grains belonging to the Ko¨selerli Formation (Gedik et al. 1979). Nannofossils are used to date the carbonate deposits in the Ermenek using the nannofossil stratigraphy of Martini & Muller (1986). Three chronostratigraphic units can be distinguished in the Neogene marine sediments. The lowermost carbonate sediments are poorly dated as the upper part of nannoplankton zone NN4 which corresponds to the latest Burdigalian and early Langhian (Berggren et al. 1995) (Figs 2 & 3). This date is speculative because it is based on one sample only. Most of the strata in the Ermenek Platform contain nannofossils from zone NN5 that correspond to the Langhian and early Serravallian. The uppermost part of the studied section has been dated, based on 3

The lowermost marine limestone interval overlies fluvial conglomerates and onlaps the pre-Miocene palaeotopography (Fig. 2b and lower 3D diagram in Fig. 3). The general depositional geometries consist of horizontal parallel bedding except in a few areas where consistent inclined bedding, ranging from 2–108, is observed (Fig. 3). Carbonate facies include oyster floatstone, associated with bioclastic packstone with common bivalves and bryozoans at the base. The upper part of this succession is dominated by red algal, bryozoan and large benthic foraminifera wackestone –packstone. Based on the depositional geometries and the facies association, this succession is interpreted as having been deposited on a marine open shelf with a main margin facing the Ermenek basin to the south (around the current Ermenek town; Figs 1c & 3). A few kilometres to the north of Ermenek town, south- and north-dipping beds indicate the presence of a seaway or intra-shelf depression (Fig. 2a and lower 3D diagram in Fig. 3). The upper limit of this interval is a clearly defined, marine flooding surface where marly planktonic-rich deposits overlie red algal, bryozoan and large benthic foraminifera wackestone – grainstone facies. This vertical facies change can be observed almost everywhere in the studied area except near the pre-Miocene substratum highs where the red algae, bryozoans and large benthic foraminifera facies are overlain by very similar facies rich in red algae and large benthic foraminifera

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(at 25 m in the sections in Fig. 3) indicating a landward retreat of the shallow-water area toward the substratum highs associated with a relative sea-level rise. Nannofossils belonging to the NN5 zone have been found within the first few metres of the marly deposits.

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Langhian: coral-rimmed platform A general palaeogeographical distinction can be made between a deep basin in the south, a carbonate-dominated shelf area in the middle, and the exposed substratum in the north (Fig. 2b and

Fig. 3. Composite sedimentary log and 3D palaeogeographical reconstruction of the platform during the late Burdigalian (bottom), Langhian (middle) and Serravallian (top).

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middle 3D diagram in Fig. 3). At the margin, between the shelf and the deeper basin to the south, a narrower vertically aggrading accumulation of shallow-water carbonate is asymmetric in shape. The basin-facing beds have steep dips (up to 508) whereas the lagoon-facing beds have dips up to 258. Regional mapping by Janson (2002) has shown that this narrow coral and red algal boundstone dominated platform forms a discontinuous barrier interrupted by narrow seaways at the shelf– slope transition. During the Langhian, the carbonate shelf consisted of a patchwork of prograding and aggrading, small, attached, carbonate platforms (Fig. 2 and middle 3D diagram in Fig. 3). The relief between the flat-bedded platform tops and the flat-bedded marly deposits between the platforms ranges from 15–35 m. The platforms show multi-directional progradation radiating from the substratum highs where they nucleated (Fig. 2a). The platforms eventually coalesced to form a single large shallowwater platform that covered almost the entire study area (Fig. 2b and middle 3D diagram in Fig. 3). By the end of the Langhian, only a narrow intra-shelf depression existed between the barrier and the coalesced platform (Janson 2002). The Langhian upper boundary is a regional exposure surface, recognizable across most of the study area (Janson 2002). Facies in the Langhian interval consist mostly of red algal, large and small benthic foraminifera and coral packstone to framestone in the platform area and marly planktonic foraminifera wackestone in the intra-shelf depression. A more detailed analysis of the facies and its distribution in the Langhian interval is described in the following sections of this paper.

Serravallian: muddy shelf with algal banks and foraminifera shoals The Serravallian interval is different both in terms of the architecture as well as the faunal content of the carbonate system. On the shelf, it is characterized by isolated algal banks and bioclastic shoals within a generally muddier environment (Fig. 2b and upper 3D diagram in Fig. 3). At the deep basin margin, a rimmed platform edge, similar to the one in the Langhian, still remained (Fig. 2b and upper 3D diagram in Fig. 3). Facies in the shoals are dominated by encrusting red algal debris and abundant rhodoliths, as well as abundant Heterostegina foraminifera, bivalves and echinoderms. The deeper areas between the higher energy shoals and banks are dominated by muddy wackestone rich in Heterostegina, bivalves, gastropods and echinoderms. Some siliciclastic deposits occur near the substratum highs in the north (Fig. 2a and upper 3D diagram in Fig. 3). These terrigenous deposits are carbonate-

cemented conglomerates that yield centimetre to millimetre, rounded substratum pebbles associated with various amounts of bivalves and other marine bioclasts (upper 3D diagram in Fig. 3). Sedimentary structures within these conglomerates consist of large unidirectional inclined foresets. Based on the facies and the sedimentary structures, this conglomerate package could be interpreted as tidal bars (Allen 1980a, b).

Stratigraphic architecture of the Langhian system Stratal geometries The studied platform margin is approximately 1.2 km long, 250 m high and continuously exposed along the sides of several canyons (Figs 2a & 4). A geometrical model was constructed at true scale based on surface tracing in the field and on photo panels, which were subsequently tied-in to the vertical measured sections (Figs 5–7). At the large scale, the geometrical evolution starts with a flat bedded aggrading ramp up to surface 2 (Fig. 6a). The horizontal beds below surface 2 dip down at a low angle outside the study area, approximately 1 km further north. Surface 2 marks a landward shift of the depositional system followed by an aggrading set of accretionary units of the platform margin up to surface 5. Between surfaces 2b and 3, this interval shows lensoid shaped accretionary units because the exposed section is closer to a strike section than a dip section. The same unit shows substantial progradation associated with the same aggradation in the adjacent valley where the section is closer to a true dip section (Janson 1997, 2002; Broucke 1998; Le Bec 1999; Janson et al. 2007). The following set of accretionary units shows progradation and aggradation between surface 5 and surface 9 with steep slopes (foreset slopes are between 15 –258). Between surface 9a and 9b, a downward step of a few metres occurs creating an onlapping wedge of sediments that corresponds to surface 9 on the bank top. Between surfaces 9 and 11, the margin shifts over 250 m towards the NW. Associated with this important progradation, the topset of the accretionary unit shows minor aggradation. Toplap occurs below surface 11, indicating a minimum of accommodation during this progradation pulse. From surface 10 to 15, the margin still progrades but with a more important aggradation component than just before (between surfaces 9 and 11). Accretionary units in this interval have a lower slope angle (between 5–158). This interval is overlain by the uppermost exposed interval that seems to be almost purely aggrading, but the geometries are hard to evaluate because they are not well exposed.

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Fig. 4. Topographic map of the outcropping margin studied in detail. The dashed grey line indicates the mean outcropping azimuth used for the reconstruction. The thick black segments show the measured sedimentary sections. Arrows indicate the progradation directions of the margin, and the associated numbers indicate the dip of the clinoform.

Within this overall aggrading and prograding margin, three types of geometries of accretionary units can be distinguished following the classification of Adams & Schlager (2000) and Quiquerez & Dromart (2006): (1) asymmetrical sigmoid; (2) symmetrical sigmoid; and (3) exponential wedge (Figs 5–7). (1) Asymmetrical sigmoid accretionary units are characterized by thicker slope deposits than platform top deposits and distal thinning-out of the foresets. (2) In symmetrical sigmoid accretionary units, the thickness difference between the slope and

platform deposits is less than in the asymmetrical sigmoidal accretionary unit since both clinothem bounding surfaces are almost parallel except where the clinothem thins-out into the distal bottomsets. (3) The exponential wedges are accretionary units that do not show any preservation of the platform top, commonly have internal toplap and have an exponential slope profile. The vertical difference between the platform top and the toe-of-slope ranges from approximately 20 –30 m during most of the progradation and aggradation to less than 10 m in the uppermost accretionary units at the top of the studied interval

Fig. 5. Photo-mosaic of the platform margin with the position of measured logs, indication of key surfaces and the two windows where the facies have been mapped in detail. Note the evolution of the platform margin break.

Fig. 6. (Modified after Janson et al. 2007) (a) Geometrical reconstruction of the prograding margin. The depositional geometries show an alternation of sigmoidal clinothems (orange) and more isopachous clinothems (green) corresponding respectively to decreasing and increasing A/S ratios. (b) Stratigraphic cross-section illustrating facies variability within the prograding margin. Triangles indicate the stratigraphic cycles. The blue boxes show the two windows where the facies have been mapped in detail. Location shown in Figures 2, 3 and 4.

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Fig. 7. (a) Close-up view of the prograding and aggrading accretionary units between surface 7 and surface 11 (Fig. 6) in window B. This outcrop photo illustrates the different geometries of accretionary units. The bottom photo has colour interpretation that outlines the individual clinoform and their geometries. (b) Outcrop picture and interpreted depositional geometries for window C between surfaces 11 and 14.

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where the progradation had filled most of the intrashelf depression and the different attached platforms started to coalesce. Periods of strongest aggradation are characterized by symmetrical sigmoid accretionary units, when space was created and sediment preserved on the inner platform. Periods of stronger progradation are dominated by asymmetrical sigmoid accretionary units, when accommodation on the platform top was reduced and sediment was transported to the margin. This implies that on the platform top the sediment production was greater than the space available. The exponential wedge morphology indicates a period of non preservation of platform top sediment. This geometry can be achieved by either eroding the platform top sediment accumulation of a complex sigmoid-oblique accretionary unit, or by sub-aerial exposure of the platform top, which reduces carbonate deposition and production to platform margin and the basinal areas. The difference between the two mechanisms can be distinguished only by toplap geometries in the case of erosion and toplap and onlap geometries onto the previous accretionary unit in the case of sub-aerial exposure. The only evidence for sub-aerial exposure is provided by the presence of Microcodium, observed on bedding surfaces in hand samples and thin section. This geometrical framework, constrained by the observed surfaces, provides the framework for the study of the facies distribution (Fig. 6b).

Facies definitions Facies models. The facies analysis is based on both macroscopic and microscopic observations. The environmental interpretation is based on the texture, the faunal composition and the position along the palaeodepositional profile. Eleven facies have been defined (Fig. 8) and illustrated in Figure 9. These have been grouped into five facies associations taking into account differences in depositional geometry and faunal content (Fig. 10). Association 1: oligophotic red algal, oysterdominated ramp to very low angle platform top. These facies occur along a flat-topped platform profile (ramp) with slope angles at the margin ranging from 5–158 (Fig. 10a). The following facies can be found along the depositional profile: oyster floatstone and framestone facies (F3, Fig. 9d), coarse-grained (F10, Fig. 9k) and fine-grained algal facies (F11, Fig. 9l), bioclastic grainstone to rudstone (F6, Fig. 9f) and large benthic foraminifera wackestone–packstone (F8, Fig. 9h, i) and planktonicdominated wackestone (F4, Fig. 9e). This facies association is dominated by oligophotic to aphotic organisms (Pomar 2001). The

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most common and extensive facies are the red algal dominated facies (F10 and F11) and the bioclastic packstone with common bivalves, bryozoans, benthic foraminifera fragments (F6). Oyster buildups a few metres high and 10– 30 m wide, scattered on the platform top are typical for this environment. Oyster debris floating in wackestone to packstone matrix is common around the buildups and as individual beds. Large benthic foraminifera consist mostly of Amphistegina and Heterostegina with few milliolids in the more landward area. The most downdip facies, in the intra-shelf basinal depression, consists of planktonic foraminifera-dominated wackestone with fragments of Heterostegina and common scaphopodia (F4). This intra-shelf basinal facies is present in all depositional models and therefore will not be repeated hereafter. Association 2: Flooded or dolomitized low angle platform margin. This facies association occurs along the margin and locally at the platform top (Fig. 10b). It is characterized by large benthic foraminifera wackestone (F8), which dominates almost the entire depositional profile in association with fabric-destructive dolo-wackestone and potentially dolo-mudstone (Fig. 9j). This facies association occurs in sigmoid clinoforms that aggrade over the platform top. It is interpreted as a flooding of the entire platform with dominance of aphotic to oligophotic biota along the entire depositional profile. The exact timing and mechanism of the dolomitization has not been studied. This facies association occurs locally, in few-metre-thick successions, at the base of the studied section (below surface 5 in Fig. 6b). Association 3: Red algal and coral dominated open platform. This facies association consists of the following facies: coarse- and fine-grained, red algal facies (F10 and F11), coral boundstone (F1), red algal bindstone (F2), large benthic foraminifera facies (F8), and planktonic-dominated wackestone (F4) (Fig. 9c). The most common facies are the red-algal-dominated facies (F10 and F11) that are present along almost the entire depositional profile (Fig. 10c). Few red algae and coral buildups can be found at the slope break and on the slope. The coral buildups are small (less than a few metres high and a few tens of metres wide). They do not form a continuous barrier that could tamper the hydrodynamic energy at the platform margin; they occur as isolated patch reefs at the platform margin. Large benthic foraminifera and scaphopod debris are the most common skeletal fragments found as the toe-of-slope biota. Further downdip the number of planktonic foraminifera increased as the matrix became muddier. This facies association is dominated by oligophotic organism with a few euphotic

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Fig. 8 Facies description.

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organisms such as coral. This facies association is the most common within the sigmoid and sigmoid oblique clinoforms between surface 2 and 10 in Figure 6b. Association 4: Coral and porcellaneous foraminifera dominated semi-restricted platform. This facies association contains the same facies as the red-algal- and coral-dominated open shelf (3) but the platform interior is dominated by small porcellaneous foraminifera muddy packstone (F7) instead of fine, red algal facies (F11) (Fig. 10d). Coral patch reefs are scattered on the platform top, at the platform margin, and on the platform slope. Compared to facies association 3, these coral buildups are much wider and slightly taller. The slope is dominated by coarse-grained red algal facies with both local and platform-derived foraminifera. Exceptionally (between surface 14 and 15) oyster floatstone facies can be found near the margin area. This association is dominated by euphotic to mesophotic organisms. Coral buildups are common but do not create a continuous rimmed margin. Nevertheless, the platform interior is muddier than in facies association 3, and is dominated by small porcellaneous benthic foraminifera typical of tropical lagoon and back-reef deposits in modern tropical systems. Alternatively, this texture change can be interpreted as the result of the stabilizing effect of sea grass meadows, similar to the Posidonia meadows in the modern-day Mediterranean (Pomar 2001). This facies association occurs mostly above surface 11 in Figure 6b. Association 5: Coarse red algal dominated platform margin wedge. This facies association is dominated by coarse, red algal facies (F10) with minor amounts of coralgal rudstone (F5) and rare small coral and red algal buildups (F1) (Fig. 10e). It occurs as oblique accretionary units that onlap on the previous clinoform. The platform top shows no sediment accumulation as a result of sub-aerial exposure or non-deposition (bypass) due to hydrodynamic energy. Foraminifera assemblages The detailed microfacies analysis of the foraminifera population, carried out on 265 samples, helped to constrain the environmental interpretation. Because foraminifera are abundantly present in all the observed facies, their distribution along the prograding margin can help to identify small-scale changes of depositional environment that might not be recognized using the general macroscopic facies classification alone. Four foraminiferal assemblages are distinguished and their palaeoenvironmental interpretation is based on their position along the depositional profile, as observed in the field, and on previous

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studies on foraminifera ecology (Gupta 1999; Hallock 1999; Geel 2000). Planktonic foraminifera assemblage: distal slope and basinal deposits. This association is characterized by up to 40% planktonic foraminifera (Globigerinidae) associated with up to 10% debris of large benthic foraminifera (mainly Heterostegina). This association is typically found in the intra-platform basinal deposits and represents the relatively deepest environment. Amphisteginidae and Numulitidae assemblage: platform margin. This association contains Amphisteginidae (Amphestegina) and Numulitidae that are represented mostly by Heterostegina and/or Operculina, which are difficult to distinguish from Heterostegina in thin section as well as in hand sample. Typically, these large benthic foraminifera represent at least 30% or more of the foraminifera content. Similar foraminifera associations are considered to be diagnostic of high-energy environments of the fore reef (Gupta 1999; Hallock 1999). In this study, Amphistegina and Numulitidae are often observed in the accretionary unit slope. This association represents an open marine environment, in a basinward position of the platform margin or close to the platform margin where hydrodynamic energy is relatively high. Porcellaneous foraminifera assemblage: internal platform. This assemblage consists of different types of small porcellaneous benthic foraminifera, such as Miliolidae (Spiroclina, Quinqueloculina and Triloculina) and Rotaliidae. Moreover, the presence of large Soritidae, Peneroplidae (Dendritina, Spirolina, Laevipeneroplis, and Peneroplis) as well as Elphidiidae (Rotaliina, Elphidium and Cibicites) and Nodosariidae are typical for the restricted environment. Presence of Peneroplids and Soritids (.5%) associated with common (.20%) Mioliolidae (Quinqueloculina and Triloculina) and Alveolinidae such as Borelis Melo is diagnostic of this association, which is typical for this proximal and protected environment observed in the platform interior area. The possible presence of sea grass in this setting would have provided a specific niche for these foraminifera. Alveolinidae and Numulitidae assemblage: Mixed proximal/distal environment. This assemblage is characterized by the presence of Alveolinidae (mostly the species Borelis Melo) and Miliolidae that are commonly associated with Amphisteginidae and Numulitidae and very rare Peneroplidae, except sparse Dendritina. This faunal composition is a mix of specimens typical of both proximal and distal environmental conditions, and represents thus a transitional environment between fully open and restricted environments.

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Fig. 9. Field pictures of (a) a massive coral framestone and (b) a branching coral framestone; (c) slab of a red algal bindstone; (d) field picture of an oyster biostrome; (e) photomicrograph of a planktonic rich wackestone; (f) slab of a bioclastic coarse-grained grainstone to rudstone;

Based on the sample location and the facies distribution (Fig. 6b) the interpreted schematic distribution of the different foraminifera assemblages along the studied section is presented in Figure 11.

Sequence subdivision Depositional cycles were defined using the stratal geometry and facies and faunal distribution. Based on the geometrical pattern, the prograding margin

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Fig. 9. (Continued) Photomicrographs of (g) small porcellaneous benthic foraminifera facies, Avl. Alveolinidae; Mil., Miliolidae; Pen., Peneroplidae; (h and i) large benthic foraminifera facies (h) Heterostegina dominated and (i) Amphestegina dominated; (j) dolomitized facies; (k) medium-grained and (l) coarse-grained red-algal-dominated grainstone/packstone.

can be subdivided in three orders of stratigraphic cycles (Janson 1997, 2002; Broucke 1998; Le Bec 1999; Janson et al. 2007): (a) at the large scale, covering the overall transgressive-regressive sequence of the late Burdigalian and Langhian interval

(NN5 nannofossil zone); (b) at the medium scale, defined as 2 sequences in Figure 6b and Figure 12; and (c) at the small scale, comprising 8 cycles represented by the smaller triangle in Figure 6b and described in Figure 12. The stratigraphic

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Fig. 10. (a)–(e) simplified depositional models for the Langhian succession. These models illustrate the lateral facies relationship for each depositional profile recognized in the studied section.

architecture of the margin will be described here concentrating on the medium-scale cycles, which allow illustrating the dynamics of the prograding margin. Higher order cyclicity was recognized in this succession but will not be described in detail here because it provides less insight into the general evolution and the controls on the dynamics of this carbonate system. The large-scale and medium-scale cycles are summarized in Figure 12. Large scale sequences. The late Burdigalian/ Langhian succession represents one overall transgressive-regressive sequence that can be subdivided into two large-scale sequences based on the depositional geometries and the vertical and lateral facies evolution. The dashed line of Figure 6a shows the evolution of the platform margin break, and emphasizes alternating periods of aggradation, aggradation and progradation, and progradation alone. The first large-scale cycle consists of an aggrading succession from the bottom of the section up to surface 2, and a more prograding and aggrading

interval from surface 2 to surface 9 or 11. Surface 2 corresponds to a regional backstep of the shallowwater carbonate facies. In the Ermenek area no clear indication of exposure has been found (Janson 2002). However, Bassant et al. (2005) documented a sequence boundary at the NN4/NN5 boundary. This surface may be both a regional exposure surface and a regional flooding surface. In that case, the lower part of the outcrop succession (NN4) should not be included in this cycle (Figs 2b & 4). In the aggrading interval, bioclasts consist of abundant small oyster buildups and oyster debris, as well as fragments of bryozoans, bivalves, large benthic foraminifera such as Heterostegina and Amphistegina and various amounts of red-algae debris (facies association 1). The turnaround point from aggradation to aggradation and progradation is characterized by tightly cemented and dolomitized rocks (facies association 2) above surface 2 (Fig. 6b). In the aggrading and prograding interval, topsets and foresets of the accretionary units are dominated by red-algae debris facies

MIOCENE PLATFORM MARGIN FROM SOUTHERN TURKEY Fig. 11. Distribution of the foraminifera associations in the studied margin. This cross-section is based on foraminifera associations identified in thin sections. The distribution is based on discrete samples but the lateral extent follows the distribution of macro-facies observed in the outcrop. This figure illustrates how the foraminifera assemblages vary with the depositional geometries. Superimposed on the two large-scale sequences, the foraminiferal content also varies within the medium- and small-scale cycles.

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Fig. 12. Description of the large-scale and medium-scale stratigraphic cycles.

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with few coral and oyster buildups, bivalves fragment, benthic foraminifera and echinoderms debris (facies association 3). The top of this first large-scale cycle can be located on surface 9 where there is a small (,2 m) downstep of the carbonate platform creating an onlapping exponential wedge (between surface 9 and 9b) with a clear toplap truncation within the wedge (Figs 6 & 7a). These geometries are the only place where an onlapping wedge with toplap is observed. Based on geometrical interpretation this would indicate a relative sea-level drop (negative accommodation). Alternatively, the top of the first cycle could be placed at surface 11 where there is a major change of fauna in the platform interior changing from fine-grained red-algal-dominated facies (F11; Associations 3 & 5) to small porcellaneous foraminifera-dominated facies (F7; Association 4) combined with an exponential accretionary wedge unit with toplap and a major seaward step of the overlying accretionary units. The faunal change is not a local shift since it can be observed on a section measured 1.5 km landward from the margin (section L in Fig. 6b) and near the nucleation point of the platform (Lebec 1999). The second cycle starts with an aggrading and prograding interval from surface 9 or 11 up to few metres below surface 12 followed by an interval of progradation and aggradation interval from surface 12 to surface 15. During this second Langhian stratigraphic cycle, the red-algal-dominated margin evolves into a platform with common coral buildups at the slope break and on the platform top associated with small porcellaneous benthic foraminifera facies in restricted areas. During the transgressive interval, the bank top is dominated by red algal facies (facies association 3). The turnaround from transgressive to regressive interval is marked by numerous coral buildups scattered on the platform top. The regressive part of this cycle is again characterized by facies association 4 with abundant coral patch reefs on the bank top and at the bank margin. Apart from one interval when red-algaldominated facies (facies association 3) recolonized the entire bank top, facies association 4 persists until the exposure horizon, with Microcodium, that terminates this cycle. Medium-scale cycles. Eight medium-scale cycles can be defined and a description of accretionary unit type, facies associations and thickness for each medium-scale cycle is given in Figure 12. Six out of 8 of these cycles are composed of a symmetrical sigmoid accretionary unit with significant aggradation on the platform top at the base and an asymmetrical sigmoid or exponential wedge with less or no platform top aggradation. These accretionary unit sets typically include two to four individual accretionary units (Figs 6 & 7). The vertical facies

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differentiation in each cycle is less pronounced at this scale, and some cycles have the same facies association in the transgressive and regressive part of the cycle, whereas in others the foraminifera association varies within a medium-scale cycle (Fig. 11). The sedimentary dynamics are generally expressed by the changes in clinoform morphology. The cross-section in Figure 11 is based on punctuated information where the samples were taken. The foraminifera assemblages were extrapolated between the samples using the geological model of Figure 6b based on observation of the macrofacies. The reconstructed distribution of the four observed foraminifera assemblages (Fig. 11) shows a similar evolution to the one recognized in the macrofacies section (Fig. 6b). During medium- and large-scale transgressive hemicycles, the external and mixed foraminifera associations tended to colonize the slope and eventually the entire platform area. During regressive hemicycles, the internal foraminifera association was present on the platform top while the mixed and external foraminifera associations moved basinward on the depositional profile. These changes are modulated by the ecological evolution of the margin. The lower part of the section is dominated by external and mixed foraminifera associations whereas in the upper part, the internal foraminifera association is conspicuous on the platform top and the medium-scale cycles are only expressed by the extent of the mixed association before or after the accretionary unit slope break. This pattern is remarkably similar to the study of an Eocene margin by Geel (2000) who proposed a similar partitioning of the foraminifera population between time of transgression where perforated large benthic foraminifera dominated the foraminifera population and regression times where Miliolidae and small imperforated foraminifera proliferated and dominated the faunal content.

Discussion Sea-level The detailed analysis of the Ermenek platform margin allowed proposing a general relative sealevel curve for this region for the late Burdigalian to early Serravallian (Fig. 13). This curve shows two large-scale sea-level fluctuations. The estimated amplitude of these fluctuations is estimated in the order of 20–40 m. This fluctuation range is based on: (1) The inferred bathymetric difference between exposure horizons with microcodium (surfaces 13 and 15 in Fig. 6) and the overlying coarse red-algal-dominated facies (Adey 1979; Rasser & Piller 1997).

284 X. JANSON ET AL. Fig. 13. Early to middle Miocene age, stages, biozonation, sequences, isotopic curves and sea-level reconstruction modified after Piller et al. (2007). Geochronology, geomagnetic polarity chrons, biozonations of planktonic foraminifers and calcareous nannoplankton, after Lourens et al. (2004). Sequence stratigraphy and sea-level curve (after Hardenbol et al. (1998) and oxygen isotope stratigraphy (dashed curve after Abreu & Haddad (1998) partly recalibrated and correlated to regional chronostratigraphy of the Central ParaTethys from Piller et al. (2007). The solid oxygen isotope curves come from Miller et al. (1998). The sea-level reconstructions come from the New Jersey continental margin (solid line from Kominz et al. (2008) and dashed line from Miller et al. (2005)). Sea-level curves for the Mut basin from Bassant et al. (2005) and from this study. The climate column contrast an oxygen isotope derived relative temperature curve from Zachos et al. (2001) with two mean annual temperature curve for Germany (Mosbrugger et al. 2005) and Turkey (Akgu¨n et al. 2007).

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(2)

The inferred bathymetric difference between the red-algal-rich facies and the large benthic foraminifera-dominated facies (Adey 1979; Boichard et al. 1987; Rasser & Piller 1997; Hallock 1999). (3) The reconstructed topographic difference between the platform top and the intra-shelf depression which is never greater than 40 m. As a result a sea-level drop of more than 40 m would bring shallow-water facies all the way down to the intra-depression floor and these facies would be interbedded with the planktonic rich marls. Superposed on this overall trend there were probably higher frequency fluctuations. When comparing our results, which occur in the upper part of nannofossil zone NN4 and the NN5 zone on the international time-scale (Lourens et al. 2004) with other records of Miocene sea-level fluctuations, there seems to be a consensus for two Langhian sequences (Fig. 13): (a) Although Hardenbol et al. (1998) had only one Langhian sequence bounded by the Burdigalian/Langhian exposure (Lan1/Bur5) and the Lan/Serr 1 sequence boundary (Fig. 11), more recent dating places the Hardenbol et al. (1998) first Serravallian sequence (between Lan2/Serr1 and Serr2 boundaries) entirely within the NN5 zone. (b) Haq et al. (1987) defined two sequences in the Langhian, the TB2.3 and TB2.4 sequences (Haq et al. 1987). (c) Similarly, in the Gulf of Mexico, early to middle Miocene deposition is recorded by two distinct third order sequences (Hentz & Zeng 2003). (d) In the ParaTethys, the Bademian, which corresponds to the Langhian and the base of the Serravallian, is composed of three units, the lower two occurring within the NN5 zone (Fig. 11). Piller et al. (2007) correlate the lower Bademian cycles with the Bur5/ Lan1 sequence of Hardenbol et al. (1998), whereas the middle Bademian unit corresponds to the lan2/Serr1 sequence. The third Bademian unit is within the NN6 zone and may correspond to the uppermost part of the section exposed in the Ermenek area (Bassant 1999; Janson 2002). (e) Finally, the most recently published global sea-level curves (Miller et al. 2005; Kominz et al. 2008) show two sequences within the latest part of NN4 and the NN5 nannofossil zone with amplitudes of sea-level changes of approximately 20 –40 m (Fig. 11). These two regionally and globally recognized sequences were deposited at times of high sea-level

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when the connections between the Mediterranean Sea, the Indo-Pacific and the ParaTethys were open. In conclusion, there is ample evidence for a global eustacy in the two Langhian sequences. The driving mechanism for the early middle Miocene sea-level fluctuations is now widely accepted to have been of glacio-eustatic origin (e.g. Miller et al. 1998, 2005; Zachos et al. 2001). Few examples of high resolution stratigraphic records for this time of the Miocene exist. Miller et al. (1998) showed the presence of several orders of cyclicity in the isotopic record and Browning et al. (2006) and Kominz et al. (2008) identified between Mi2 and Mi3 approximately 6–10 higher frequency cycles in the North American continental shelf record (Fig. 11). In the Gulf of Mexico, Hentz & Zeng (2003) recognized 8–10 fourth-order sequences within the Langhian. These studies suggested that during this time period a 400 ka cycle dominated the oxygen isotopic record and thus the eustatic sea-level fluctuations (Miller et al. 1998). The eight medium-scale cycles recognized in the Langhian prograding margin in Ermenek match this North American record fairly well, supporting a global control of these sequences. Taking into account the estimated duration for the Langhian of 2.32 Ma, the eight interpreted sequences would have a duration of less than the proposed 400 ka if they were all strictly within the Langhian. However, the limited amount of biostratigraphic data does not enable the Burdigalian/ Langhian boundary nor the Langhian/ Serravallian boundary to be identified precisely. Alternative explanations could be that our eight sequences represent 100 ka cycles and that we record only one out of three on the Ermenek platform, or that some of the geometrical changes in the clinoform interpreted as sequences are produced purely by sediment accumulation variation rather than by change in accommodation.

Large-scale platform evolution: linked accommodation and sediment production The amount of carbonate sediment produced at any given point along a depositional profile depends on a variety of environmental conditions such as light, hydrodynamic regime, antecedent topography and trophic level. The amount of carbonate deposited at any given point depends on the in-situ production and the balance between the import and export of sediment at this location, which in turn depends on the slope angle, the grain size, shape and bulk density and the hydrodynamic regime (Pomar 2001). Modern red algae can exist from the upper shoreface environment down to more than 200 m of water depth with various morphological forms such as branching, encrusting or rodolith forms

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(Liddell et al. 1997). Similarly, the bathymetric and ecological distribution of large benthic foraminifera is quite large. Heterostegina can be found in water ranging from 20 –500 m deep (Hallock 1999). The common small benthic foraminifera identified in this study also have a widespread ecological and bathymetric distribution. Because the ecological distribution and bathymetric range of all these organisms overlap, the exact reconstruction of the palaeo-environmental condition of these Miocene deposits is problematic. The gradual change of the faunal assemblages observed in this study can be explained either by a climatic control or by the variation in water depth and associated changes in the light penetration. At the base of the section (Fig. 6b), the carbonate organisms are typically oligophotic (facies association 1) that currently thrive in the Mediterranean at depths of more than 50 m and accumulate down to 120 m (Liddell et al. 1997). The faunal assemblage changes gradually to include more mesophotic organisms with some coral buildups (facies association 2) up to surface 10. Above surface 10, small benthic porcellaneous foraminifera and small coral buildups, which are euphotic to mesophotic organisms, dominate the platform top whereas the oligophotic organisms such as red algae and large benthic foraminifera dominated the clinothem. Between surface 13 and 14, oligophotic organisms dominate the platform slope and the platform top before the system switches back to mesophotic to euphotic condition between surface 14 and surface 15. Light conditions on the seafloor depend on both water depth and water turbidity. Therefore the overall evolution of the margin could be seen as a shallowing-upward or cleaning-upward trend from the bottom of the section to surface 13 with a slight deepening between surface 13 and 14 and a second shallowing upward between surface 14 and 15. Alternatively, the change from the red algae, large benthic foraminifera and bivalves at the base to a system dominated by red algae, large and small benthic foraminifera and coral at the top can be related to a temperature control linked to large-scale climatic change (Pomar & Hallock 2007; Piller et al. 2007). The climatic evolution for the Neogene shows a relatively cold period during early Miocene followed by warmer temperatures during the late Burdigalian. Climate reached an optimum in the Langhian followed by another cooling during the middle Serravallian (Woodruff et al. 1981; Roegl & Steininger 1984; Sun & Esteban 1994; Gebhardt 1999; Abreu et al. 2000). The studied margin seems to follow these climate changes (Fig. 13): from a warm-temperate rhodalgal flat-topped shelf, dominated by oligophotic to aphotic carbonate producing organisms in the

relatively cooler late Burdigalian, to a sub-tropical to tropical platform in the warmest Langhian dominated by mesophotic to euphotic organisms. The change from mesophotic to euphotic organisms (facies association 4 and 5) to oligophotic organisms dominated Serravallian facies association with abundant red algae, large benthic foraminifera and mollusc coincides with the climate cooling of Serravallian times. In addition, in the Mediterranean area, coral diversity peaks in the Burdigalian and then dropped during the Langhian and Serravallian. Pomar & Hallock (2007) hypothesize that the high temperatures during the Langhian are not favourable for extensive coral growth. In addition, several studies have associated the red algal bloom during Miocene times with an increase of terrestrial runoff in the Mediterranean associated with the middle Miocene climatic optimum (Halfar et al. 2000; John et al. 2003). Determining whether these ecological changes are related to temperature, aridity or weathering requires further detailed study (John et al. 2003; Akgu¨n et al. 2007). At the small scale cycle, there seems to be a weak vertical facies differentiation between the increasing accommodation and decreasing accommodation part of each medium-scale cycle. This is mostly expressed in the benthic foraminifera composition of the platform top and margin area facies. Quiquerez & Dromart (2006) show that the vertical facies differentiation in clinoforms is expressed better at the base of the section compared to the upper part. Even with detailed quantification of benthic foraminifera population within a few cycles, no systematic vertical evolution of the biological component within each medium-scale cycle could be related to a climatic forcing.

Comparison with other Miocene platforms in the Mediterranean Pomar & Kendall (2008) compared several carbonate platforms, of which four examples are Mediterranean Miocene platforms, including the Ermenek margin. Pomar & Kendall (2008) compared these platforms in terms of platform types, sediment composition, stratigraphic architecture and controls on carbonate deposition. They concluded that the main difference between these four Miocene platforms related to sediment production and depositional environments combined with the hydrodynamic setting and the antecedent topography. The two Turkish early to middle Miocene platforms mentioned in this study, one near Mut (Bassant et al. 2005) and one near Ermenek (this study), are both dominated by oligophotic organisms and resulted in flat-topped open platforms similar to the present day Far Eastern platforms (Wilson

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& Vecsei 2005). These types of platform record large- and medium-scale stratigraphic variations (third-, fourth- and possibly fifth-order). In contrast, the late Miocene rimmed platform in Majorca is dominated by a euphotic, coral-dominated carbonate factory similar to the present day Caribbean and Indo-Pacific platform. This type of platform is very sensitive to changing accommodation and therefore records a higher level of cycle (third- to sixth-order). Distally-steepening ramps are dominated by oligophotic and/or aphotic carbonate-producing organisms and are less sensitive to change in accommodation (second- and third-order only). This type of platform generally records large-scale changes in hydrodynamic setting, biota evolution linked to climate control rather than changes in hydrodynamic regime (Pomar et al. pers. comm.). When comparing the Langhian Ermenek platform with the Burdigalian Mut platform, they share similar facies and facies associations, typical of an oligophotic to mesophotic carbonate system, and eustatic amplitudes were in the same order of magnitude (Bassant et al. 2005; Kominz et al. 2008), possibly the only difference being that the climate in the Burdigalian was potentially slightly cooler and more arid (John et al. 2003; Akgu¨n et al. 2007). However, they differ in the antecedent topography, which influenced their dimensions and architectures. In the Mut area, attached carbonates platforms, and mixed carbonates/siliciclastic shelves developed around the margins of a deep basin during transgressive systems tracts (TSTs) and highstand systems tracts (HSTs), whereas small isolated platforms developed on top of older delta fronts during lowstand systems tract (LSTs) which subsequently drowned during TSTs (Bassant 1999, 2005). In the Ermenek area, the antecedent topography consisted of intra-mountain depressions filled with lacustrine and fluvial facies (Janson 2002) with no significant deltas that could promote the development of a mixed shelf and small isolated platforms. In contrast, small attached platforms nucleated on basement highs and aggraded and prograded into relatively shallow water while an aggrading rimmed platform developed at the shelf margin. The horst-related topography in the Ermenek area created the ideal substratum for the development of a 20 km large carbonate shelf where water depth never exceeded more than 50 –75 m (Janson 2002).

Conclusions The exceptional outcrop conditions of the late Burdigalian, Langhian, and early Serravallian carbonate system has permitted detailed study of the

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relationship between depositional geometries and sedimentary facies in this time interval. (1)

(2)

(3)

(4)

A general evolution of the sedimentary system is observed from a late Burdigalian/early Langhian open shelf, to a Langhian, coralrimmed amalgamated platform system, to a Serravallian muddy shelf with algal banks and foraminiferal shoals. In the Langhian system, which was the focus of this study, a clear hierarchical stratigraphic organization with three orders of sequences was distinguished: an overall transgressiveregressive sequence covering the Langhian stage, which comprises 2 large-scale sequences and 8 medium-scale cycles. These have been defined by the very well-expressed depositional geometries and the faunal content. The large scale Langhian sequence shows an evolution from aphotic and oligophotic faunal assemblages in the lower part of the sequence, associated with mostly aggrading and slightly prograding geometries, whereas the facies in the upper part of the sequence are characterized by a mesophotic to euphotic faunal assemblages. This evolution is observed along the entire platform top profile from the margin to the coastal onlap. The medium-scale cycles are expressed by changes in the clinoform geometries, whereas facies changes are more subtle and take place over several cycles. Three controlling factors influenced the depositional patterns of this sedimentary system: (a) Sea-level fluctuations Both the overall transgressive-regressive Langhian sequence, and the 2 large-scale sequences have been recorded worldwide, and are probably of glacio-eustatic origin. The medium-scale sequences cannot be linked directly to other recorded sea-level fluctuations. (b) Climate The overall biotic change observed both along the studied margin, and regionally (Janson 2002), is in phase with the climatic changes observed during this time: a relatively cool Burdigalian, followed by the mid Miocene climate optimum during the Langhian, followed by another relative cooling in the Serravallian. A response of the faunal assemblage to the changing climatic conditions is our preferred interpretation for the biotic change. (c) Antecedent topography This is a local controlling factor that explains the difference in architecture between the Ermenek prograding, coalesced platform

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system which developed on a relatively shallow shelf in the western Mut basin, compared to the land-attached and isolated platforms, and mixed shelves that developed in the much deeper, and riverfed eastern Mut basin. There is a clear global climatic and eustatic signal expressed in this early to middle Miocene carbonate system, which controlled the stratigraphic architecture and the dominant carbonate producers (faunal assemblages). The local expression was further controlled by antecedent palaeotopography, providing a substratum for the development of a carbonate shelf. These exceptional exposures have a potential for further study of the dynamics of shallow-water carbonate systems to large-scale climatic change and eustatic sea-level fluctuations. Between 1996 and 1999 a multi party research team led by the IFP, in cooperation with Total and GDF, worked at several locations in this study area. Masters theses by O. Broucke, X. Janson, A. Lebec and A. Wattine have provided supporting material for the present paper. The assistance of J. P. Leduc and F. Cauchet with the measuring of the outcrop locations is acknowledged. Prof. N. Gorur, then of the Istanbul Technical University, provided logistical support for this project. Constructive comments by reviewers P. Bassant and G. Warrlich have helped to improve the manuscript.

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Meteoric diagenesis in carbonates below karst unconformities: heterogeneity and control factors O. WEIDLICH Institut fu¨r Angewandte Geowissenschaften, Technische Universita¨t Berlin, Ernst-Reuter-Platz 1, D-10587 Berlin, Germany Present address: Wintershall Holding AG, Friedrich-Ebert-Strasse 160, D-34119 Kassel, Germany (e-mail: [email protected]; [email protected]) Abstract: The combination of sedimentological and diagenetic data is important for the characterization of carbonate pore systems. This is particularly true for carbonates that were affected by meteoric diagenesis during sub-aerial exposure, for instance at sea-level lowstands. This diagenetic environment is commonly believed to be associated with increases in porosity, permeability and pore-throat diameters. Using data from three localities, improvement or deterioration of reservoir parameters below karst unconformities were analysed with a three-fold approach. In the first step, meteoric dissolution was characterized and early to late diagenetic products were described. In the second step, sedimentological and diagenetic data were converted to petrophysical data. In the third step, modelled climate data, in particular the occurrence of monsoon cells, in conjunction with other control mechanisms, were considered to understand the processes that controlled meteoric dissolution and later pore infill. Three case studies were analysed: (1) Lower Triassic oolites (sedimentary rocks dominated by ooids) and microbialites of the Calvo¨rde Formation (Buntsandstein Group, Germany); (2) stacked shallowing-upward cycles of carbonate platform deposits in the Middle– Upper Triassic Mahil Formation (Arabian plate, Oman), capped by palaeosols; and (3) an Upper Triassic coral patch reef and overlying strata (Adnet, Salzburg region, Austria). Data integration allowed the establishment of three scenarios of significantly different processes related to meteoric diagenesis below unconformities: (1) increase of porosity and permeability and their preservation through time; (2) increase of porosity and permeability and subsequent pore system occlusion; and (3) decrease of porosity and permeability and creation of a barrier for pore fluids. Knowledge of the time span involved in meteoric diagenesis and the nature of the climatic regime helped to explain the origins and control mechanisms of the meteoric pore systems. The study provided evidence that a wellconnected, large karst system, typical of a humid climate, is likely to be sealed subsequently by sediment and cement. Under arid climatic conditions, tight palaeosols developed at the unconformity and small karst pore systems developed which had the potential to remain open during basin evolution. Depending on the aforementioned parameters, carbonates affected by meteoric diagenesis may either become tight rocks or reservoirs.

Prediction of diagenetic products and quantitatively defining the parameters controlling porosity and permeability development at outcrop scale are important for reservoir rock typing and static reservoir modelling in the oil industry. Diagenetic processes, especially meteoric dissolution and cementation phenomena, are among the most challenging input parameters of reservoir models. Some of these products, which formed along sub-aerial unconformities during eogenesis, were controlled by the amplitude of sea-level changes and the time involved in the formation of the unconformity. A variety of additional factors control diagenetic pathways of these carbonates, notably palaeolatitude, oceanography, basin configuration and depositional profile. In humid and warm areas, for example, diagenetic alterations and carbonate dissolution are quantitatively more significant than in arid and

warm areas (James & Choquette 1990b). Also, the position of the water table is of importance for limestone diagenesis because diagenetic products of the meteoric-vadose zone (pores filled with water and gas) differ from the meteoric-phreatic zone (pores completely filled with water). As a consequence, variable pore systems resulted from meteoric diagenesis with changing permeability trends. Depending on the prevailing processes during and after meteoric diagenesis, resulting products of carbonates will vary between seals and reservoir rocks. The knowledge of diagenetic products of karst unconformities has increased drastically since the books edited by James & Choquette (1988), Horbury & Robinson (1993) and Budd et al. (1995). These studies have been completed by more recent investigations of karst topography (Purdy & Waltham 1999), palaeocave carbonate reservoirs

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 291–315. DOI: 10.1144/SP329.12 0305-8719/10/$15.00 # The Geological Society of London 2010.

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(Loucks 1999; Loucks et al. 2004), time control during meteoric diagenesis (Saller et al. 1999) and the imaging of karst surfaces on seismic data (Vahrenkamp et al. 2004). Meteoric diagenesis is a well-known phenomenon of Triassic platform carbonates. Products of meteoric diagenesis have been reported from many environments of Tethyan carbonate platforms, including stacked Lofer cyclothems of the Dachstein platform (Satterley 1996; Enos & Samankassou 1998) and reefs (Satterley et al. 1994; Schroeder & Reinhold 1997 and further references herein; Bernecker et al. 1999). Subaerially exposed cycle tops were also described from isolated platforms from the Dolomites (e.g. Goldhammer et al. 1990; Egenhoff et al. 1999). This paper also focuses on meteoric diagenesis in Triassic carbonates. The aim is to study the impact of climate on meteoric diagenesis and to characterize early- to late-stage diagenetic products filling the pore space using data from three different platform environments, These factors control porosity and permeability and have a major impact on petrophysical properties of flow units. Key aspects of this analysis include: † determination of facies associations of host rock † penetration depth of meteoric diagenesis † nature and dimension of the resulting pore system (enlargement vs rearrangement of porosity) † subsequent to late-stage diagenetic and sedimentary sequences of the pore system † importance of stratabound and non-stratabound joints for the karst system † transformation of all geological data to petrophysical data. Applying the three-fold classification of palaeokarst (James & Choquette 1988), the case studies characterize phenomena either of local palaeokarst or depositional palaeokarst. Low-amplitude Triassic sea-level changes do not cause interregional karst unconformities because regional tectonic movements are sufficient to overprint high amplitude global signals of sea-level changes.

Methodology In order to gain comparable data sets for theses case studies, meteoric diagenetic products were analysed using the same procedure. All case studies were characterized with data of comparable dimensions. From outcrop to pore-scale, the following methodology was applied: † tracing of unconformities and characterization of the pore system in the field using phototransects;

† establishment of diagenetic sequences of the pore system using detailed field photographs and slabs; † quantification of open porosity and calculation of permeability from representative thin sections with digital image analysis; † petrographic description of diagenetic sequences, including SEM; and † geochemical differentiation of marine and meteoric diagenetic environments using stable isotope ratios (d13Ccarb and d18Ocarb). From about 250 polished hand specimens 50 representative samples were chosen for petrography and stable isotope analysis. Thin sections of the samples were checked prior to the start of geochemical analysis to guarantee that areas with the least overprint (e.g. recrystallization phenomena) were used for stable isotope analysis. For this purpose, the ‘white card technique’ described by Dravis (1991) was used to check whether or not cements or grains had suffered from neomorphism. The simple and cheap technique comprised a white sheet of paper which was put under the thin section during observation with plane-polarized light. The resulting diffuse light highlighted ghost structures, whereas conventional plane-polarized light emphasizes details of crystals after neomorphism. In this paper, cements were consequently regarded as pristine if the white card technique and conventional plane-polarized microscopy yielded similar results and showed no evidence of ghost structures. A dental drill was used to gain powder for stable isotope analysis from skeletal grains, sediments and cements. No bulk samples were used for interpretation. 50 samples with an average weight of 20 micrograms were analysed on a Finnigan mass spectrometer (Leibniz Laboratory, University of Kiel, Germany and Department of Geosciences, Royal Holloway, University of London, United Kingdom). All stable isotope results were reported in ‰ relative to the VPDB standard. Carbon and oxygen stable isotope data were compared with isotopic composition of ambient Triassic sea water (Veizer et al. 1997; Korte et al. 2005). The interpretation of stable isotopes in terms of the diagenetic environments is based on Veizer (1992), Veizer et al. (1997) and Bickert (2000). The standard porosity classification for carbonates (Choquette & Pray 1970), the pore size classification of Luo & Machel (1995: microporosity ,1 mm, mesoporosity 1 mm– 0.5 mm, macroporosity 0.5– 256 mm, megaporosity .256 mm) and the concept of stratabound and non-stratabound joints/ fractures (Odling et al. 1999) were used to describe the morphological parameters of pore space. Since conventional porosity and permeability data from

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plugs did not exist, porosity was quantified from thin sections impregnated with blue-dyed epoxy with digital image analysis. Using the equation of Mowers & Budd (1996), permeability was calculated from image analysis data. The results were compared to petrophysical data using the rock fabric/petrophysical classification of Lucia (1995), a modification of the classical Archie classification (Archie 1952).

Geological setting of the study areas

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ramp, see Figures 3– 6. Mixed carbonatesiliciclastic oolites (sediments rich in ooids) and microbialite levels alternate with siliciclastic fluvial and floodplain sediments. Whether these carbonates were deposited in marine or lacustrine environments has been a matter of long-standing debate (Paul & Peryt 2000). Weidlich (2007) provided new evidence for the marine origin of the oolites and microbialites. Oolites and microbialites were investigated at Harliberg (location details were described by Paul & Peryt 2000).

Three case studies representing different Triassic depositional environments were studied (Figs 1 & 2).

Middle Triassic Mahil Formation, Akhdar Group (Sultanate of Oman)

Lower Triassic Rogenstein, Calvo¨rde Formation, Buntsandstein Group (Germany)

Shallowing-upward sequences of carbonates of the Permian –Triassic Akhdar Group were deposited on the epeiric platform that rimmed the Arabian plate, see Figures 7 and 8. The Mahil Formation started with a reddish palaeosol overlying an unconformity which cut into the underlying Lower Triassic Saiq Formation. The basal Mahil Formation was studied in Wadi Aday, Saih Hatat. Location details were described by Weidlich & Bernecker (2007).

The Rogenstein was deposited in the Germanic Triassic land-locked basin along a homoclinal

Upper Rhaetian coral limestone, Tropfbruch near Adnet (Austria) The famous Adnet reef is part of the epeiric Dachstein platform, see Figures 9–12. The position was very close to the intra-platform Ko¨ssen basin. The Upper Rhaetian Tropfbruch outcrops consists of sawed quarry walls which provided a unique opportunity to investigate the ecological patterns of Mesozoic scleractinian coral communities (Scha¨fer 1979; Bernecker et al. 1999; Flu¨gel 2002). Three reef stages were separated by three karst unconformities. See Bernecker et al. (1999) for location details.

Results Lower Triassic Rogenstein, Calvo¨rde Formation, Buntsandstein Group, Germany

Fig. 1. Chronostratigraphy of the case studies. (1) Lower Triassic Rogenstein, Calvo¨rde Formation, Buntsandstein Group, locality Harliberg, Germany. (2) Middle Triassic Mahil Formation, locality Wadi Aday, Saih Hatat, Sultanate of Oman. (3) Upper Rhaetian coral limestone, locality Tropfbruch near Adnet, Austria. Time-scale in million years after Gradstein et al. (2004).

Facies association. Mixed carbonate-siliciclastic sediments of the Calvo¨rde Formation formed stacked fining-upward cycles (Hauschke et al. 1998; Szurlies et al. 1998). The ideal cycle consisted of a more or less gradual transition from sandstone to carbonate and finally to siltstone (Paul & Peryt 2000). In most cases the carbonate-rich units were either oolites (carbonate rock dominated by ooid) or microbialites (Fig. 3). Oolite textures range from packstone to float/rudstone, locally with strong bimodality in ooid size (Figs 4 & 5). Ooid types comprised normal ooids, regenerated ooid fragments, superficially incrusted ooid clasts and

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Fig. 2. Platform settings of locations (black arrows). (1) Lower Triassic Rogenstein, Calvo¨rde Formation, Buntsandstein Group (Germany); cross section through the carbonate-rich part of the Germanic basin; (2) Middle Triassic Mahil Formation, Akhdar Group (Sultanate of Oman); (3) Upper Triassic coral limestone, Tropfbruch near Adnet (Austria).

cerebroid ooids. Cerebroid ooids have an indented periphery, cortices may have tangential and radial microstructures in the same ooid (Flu¨gel 2004). The matrix of oolite contained varying percentages of detrital quartz, feldspar, mica and carbonate cement. The pore systems of the carbonate beds were enhanced or occluded by diagenetic processes described below.

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(3) Diagenetic sequence. Based on petrographic observations the following simplified sequence of diagenetic events was established: (1) Ooid dissolution At a very early stage ooids were partially dissolved in some beds during or shortly after deposition (Fig. 6a). SEM photographs and petrography documented that oomouldic porosity is exclusively separate vuggy porosity (sensu Lucia 1995), a type which did not contribute to an increase of permeability (Fig. 6a, b). The ooid dissolution affected only some beds and is not a widespread phenomenon.

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Pore occlusion by fibrous to bladed cement The first visible cement frequently observed is a thin rim of isopachous fibrous to bladed calcite rimming interparticle pores (Fig. 5d) and dissolved ooids. The thickness of this cement does not exceed 0.1 mm and tends to be thicker in partially dissolved ooids where two generations can be differentiated. Physical compaction Cementation by the aforementioned diagenetic product was minor and, as a consequence, physical compaction of the sediment led to random pressure dissolution seams along grain contacts (Fig. 5d). Microstylolites exhibit no preferred direction, they occur along vertical and horizontal grain contacts. Pressure solution affected the rocks to a varying degree. End-members can be significantly affected beds or rock units without any evidence of pressure solution. Carbonate dissolution The most obvious diagenetic product is carbonate dissolution leading to the formation of vuggy porosity. This second phase of dissolution was still

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Fig. 3. Field photograph of the Rogenstein oolite, Calvo¨rde Formation, Buntsandstein Group (Germany), illustrating sandstone-carbonate-siltstone cycles. Dashed lines highlight boundaries of oolites (sedimentary rocks dominated by ooids). Locality is Harliberg near Vienenburg (Germany). (a) Lower section with two thick oolites. Stratabound fractures of the oolite are marked with blue arrows. See hammer for scale. (b) Upper section with four oolite beds. Scale (black bar) is 2 m. (c) Close-up view of bedding plane of oolite with small microbialite on top of tilted bedding plane. Lens cap is 5 cm in diameter. (d) Detail of an oolite bedding plane. Mud cracks indicate sub-aerial exposure (white arrows). Diameter of coin 2.5 cm. (e) Close-up view (perpendicular to bedding), showing differential carbonate cementation and dissolution. Porous oolite contrasts with cemented microbialite. Diameter of cap lens 5 cm.

very early because selected open pores were filled with silt which infiltrated the pore space from overlying beds (Fig. 4a). Carbonate dissolution affected grains and fibrous to bladed

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calcite rimming the grains and, therefore, postdates the dissolution phase (1). Partial pore occlusion The dominant cement occluding open interparticle and vuggy pores

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Fig. 4. Photomicrographs of the Rogenstein oolite, Calvo¨rde Formation, Buntsandstein Group (Germany). (a) Overview of thin section with microbialite and oolite. Interparticle and vuggy porosity (black areas in thin sections) are restricted to the oolite; the microbialite is characterized by a dense fabric. Cross-polarized light. White rectangle refers to position of Figure 4b. Scale in 1 mm units. (b) Close-up view of an open vug, interpreted as meteoric-vadose diagenetic product. Note corroded ooids at the top (green arrows) and isopachous cement crust at the bottom of vug (red arrows). Plane-polarized light. Scale in 0.1 mm units. (c– d) Photomicrograph (plane- and cross-polarized light) of blocky calcite filling interparticle porosity. Scale in 0.1 mm units. (e–f) Detail (plane- and cross-polarized light) of ooid corrosion and subsequent cementation of pore space by blocky calcite. Scale in 0.1 mm units.

is blocky calcite. There is evidence that blocky calcite increased in size towards the centre of large pores (see Fig. 4b– f). The crystals attain a maximum size of 5 mm. The youngest

observed diagenetic product is idiomorphic gypsum which was observed in one sample only. Vugs and mouldic pores were not completely occluded by the aforementioned cements.

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Fig. 5. Photomicrographs of the Rogenstein oolite, Calvo¨rde Formation, Buntsandstein Group (Germany). (a) Overview of thin section under cross-polarized light, showing patchy distribution of open interparticle and vuggy porosity (black areas). Note open fractures in the lower part of the sample. White rectangle refers to position of Figure 5b–c. Width of thin section is 7.5 cm. (b – c) Photomicrograph of open pore space (plane- and cross-polarized light) which resulted from combined interparticle and vuggy porosity. Scale in 0.1 mm units. (d) Diagenetic sequence (see text for complete sequence): (2) Finely crystalline, bladed to fibrous isopachous cement rimming ooids, interpreted to be of marine-phreatic origin. (3) Pressure solution at grain contacts. (4) Dissolution of isopachous cement and ooids, probably in a meteoric-vadose environment. (5) Blocky calcite cement, meteoric-phreatic to burial diagenesis. Scale in 1.0 mm units.

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Fracturing of rocks Sandstones, siltstones and carbonate were affected by fractures (the fracture pattern was not analysed in detail). The variation in the density of fractures was controlled by lithology, for example, siltstone, sandstone or carbonates exhibit different density patterns. This phenomenon was most evident in the field. Thick carbonate beds tended to develop evenly spaced stratabound fractures which enhance the connectivity of interparticle and vuggy porosity (Fig. 3a), whereas claystone and silty sandstone yield less obvious, denser fractures. The fractures affected the rocks after partial cementation of pores. Fractures in carbonates were partially cemented by blocky calcite or remained open (Fig. 5a).

Interpretation. The thin isopachous fibrous to bladed cement rimming grains, product 2 in the diagenetic sequence, was interpreted as marine-phreatic based on morphologic criteria. James & Ginsburg (1979) described isopachous ‘stubby crusts’ from

the seaward margin of Belize, interpreted as marinephreatic, which resembles product 2 of the described Rogenstein sequence. The observed rims were too thin to deliver drill powder for stable isotope analysis using the available tools. Partial dissolution of metastable ooid layers, product 1, is interpreted to be a product of the marine-phreatic environment after consideration of the above data. Dissolution of metastable mineralogies (e.g. aragonite) was also described as product of marine-phreatic seafloor diagenesis (James & Choquette 1990a). Mud cracks at the top of the carbonate beds (Fig. 3c–d) indicated sub-aerial exposure (Paul & Peryt 2000) and provided field evidence for the meteoric alteration of the investigated beds. The stable isotope data of about 50% of the measured ooids fall in the field of marine carbonates (Veizer et al. 1997). The other ooid data (d13C ¼ 21.44 to 0.25‰; d18O ¼ 26.79 to 25.13‰) show an overlap with the field of blocky calcite, product 5 of the simplified diagenetic sequence (d13C ¼ 22.20 to 20.30‰; d18O ¼ 26.41 to 25.35‰) (see Fig. 13,

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Fig. 6. Photograph of slab and SEM photographs, showing open porosity of the Rogenstein oolite, Calvo¨rde Formation, Buntsandstein Group (Germany). (a) Photograph of slab of oolite, showing partly dissolved ooids (dark cores). Scale is 2 cm. (b) SEM photograph of mouldic porosity of a dissolved ooid. Note dense matrix. (c) SEM photograph of microporosity of matrix. Note connected pores. (d) Close-up view of well-connected microporosity of matrix.

Table 1). The relative depletion of 13C of the blocky calcite and ooids compared with the marine signal confirmed a meteoric diagenetic environment during cement formation and, thus, a meteoric origin for the pore space. The strong variation of carbon of microbialites (d13C ¼ 24.70 to 0.25‰) may reflect metabolic effects of the organosedimentary deposits. The low d18O data can be explained by burial diagenesis, a common phenomenon of Mesozoic and Palaeozoic rocks (Veizer et al. 1997; Immenhauser et al. 2002). Neomorphism (replacement and recrystallization of a precursor mineral) is a well-known phenomenon which affected microbialites (sparitization described by Paul & Peryt 2000) and cerebroid ooids (Flu¨gel 2004) to some extent. Recrystallization of pore-filling cements was not observed (see methodology section for the description of technique used).

All the carbonates studied, including beds with a thickness of 1 m, developed a range of pore systems. At one end, oomouldic porosity, interpreted as marine-phreatic, was developed. Oomouldic porosity was observed predominantly as partial ooid dissolution, affecting the nucleus plus inner laminae, and to a lesser extent as complete grain dissolution (Fig. 6a). Applying the classification of Luo & Machel (1995) the size of oomoulds falls in the range of mesoporosity. At the other end of the range, pore systems of many oolites were dominated by fabric-selective interparticle and not fabric-selective vug porosity. The size of open pores spans the range of meso- and macroporosity (Figs 3– 5). Comparison of field observations suggested that oolites and microbialites within the same bed were altered by meteoric dissolution in different ways (Figs 3d & 4). Microbialites appear to be well cemented and show little evidence of

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Fig. 7. Field photograph of dolomitized carbonates and palaeosol, Middle Triassic Mahil Formation, Wadi Aday (Oman). Some of the stratabound fractures have been highlighted with white lines. (a) Overview illustrating the transition from Lower Triassic cementstone to Middle Triassic carbonate cycles topped by sub-aerial unconformities. Note top of thick palaeosol (dashed line). Width of section approximately 100 m. (b) Detail of palaeosol deeply penetrating Lower Triassic cementstone. Note changes in the open fracture density from massive– thick-bedded dolomite. Open fractures disappear in the palaeosol.

carbonate dissolution (Figs 3e & 4a), whereas oolites yield pore systems with a mixture of interparticle pores and vugs (Figs 3e, 5 & 6). Meteoric diagenesis enhanced via preferential dissolution the existing pore systems of oolites whereas microbialites left unaltered (Fig. 3a). SEM investigations demonstrated that the carbonate matrix yielded microporosity to a varying degree (Fig. 6c–d). The timing of this type of porosity remains uncertain.

Middle Triassic Mahil Formation, Wadi Aday, Sultanate of Oman Facies association. The Mahil Formation represents the shallow-water carbonate succession deposited on the epeiric carbonate platform rimming the Arabian plate. In the study area, the Mahil Formation comprises stacked, dolomitized, shallowing-upward cycles (Weidlich & Bernecker 2003). Details of carbonate texture and sedimentary structures have been

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Fig. 8. Close-up views of palaeosol and overlying shallowing upward cycle, base of Middle Triassic Mahil Formation, Wadi Aday (Oman). (a) Mud-rich basal palaeosol with caliche hardpan. Note irregular top of palaeosol. (b) Detail of transition from palaeosol to dolo-grainstone. Sequence from base to top: reddish palaeosol, light grey grainstone and laminated dark grey wackestone. (c) Polished slab of brecciated palaeosol, showing cemented cracks (diagenetic product 3) and sediment-filled cracks and moulds of evaporites (4). Scale in mm units. (d) Thin section of sample from (c) illustrating cemented (3) and sediment-filled cracks (4). No open porosity exists. Scale in 0.1 mm units. (e) Photomicrograph of overlying dolo-grainstone (top of sample towards left). Open porosity (white areas, 2, 5) are combinations of interparticle and vuggy porosity. Dark dots are air-filled bubbles. Note iron-stained microstylolites (7). Scale in 0.1 mm units. (f– g) Close-up of dolo-grainstone illustrating ghost structures of grains (red arrow refers to the same grain). (1) micrite envelope, (2) relics of blocky calcite, (3) compaction of grains, (4) dolomitization. (f) Plain-polarized light. (g) Photomicrograph taken with white paper under the thin section. Scale in 0.1 mm units.

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Fig. 9. Overview of Upper Rhaetian coral limestone, Tropfbruch near Adnet (Austria), showing the unconformities and the position of samples. Unconformity A is penetrated by boring bivalves which colonized the hard substrate after renewed flooding.

preserved despite pervasive dolomitization. Thus, ghosts of grains and sediment structures allowed the recognition of grain-support textures at the base and mud-rich textures at the top of cycles. Shallowing-upward carbonate cycles formed sets of beds with grain- or mud-support textures and were capped by reddish palaeosols (Fig. 7a –b). For analysis, a representative grainstone overlying a palaeosol, an intraformational palaeosol, a few centimetres thick, (for stable isotope analysis) and an extra-thick palaeosol (thickness is 1.0 m or less) at the base of the Mahil Formation were chosen. The top of the palaeosols of the Mahil Formation exhibit sharp contacts, whereas there is a gradual transition to the underlying host rock at the base. The macroscopic features of sub-aerial exposure visible in the field comprised carbonate dissolution and in-situ brecciation of the host rock, reddish stain of the palaeosol and pedogenic micrite and infiltrated sediment between sheet cracks. Within the palaeosols, two facies types were distinguished, notably the caliche hard pan at the top and brecciated host rock at the bottom. Polished slabs and thin sections exhibited more characteristic features of the pedogenic carbonates. The indistinctly laminated caliche hardpan consisted of sub-horizontal micritic and microsparitic laminae, locally clotted

fabric and sediment filling the casts of roots. The brecciated dolomite consisted of clasts floating in a reddish matrix. The breccia shows all features of a reworked caliche hardpan, including multiple fracturing, iron staining of matrix and roots (Budd et al. 2002; Kosir 2004). Circular and elliptical structures, interpreted as casts of roots, were either cemented by blocky calcite or filled with red-stained micrite. Root tubules are a few millimetres in diameter. Most obvious in the breccias were multiple vertical and sub-horizontal sheet cracks and the formation of clasts. Clasts of horst rock had a varying size, ranging from few a few millimetres to more than 20 cm. The fabrics of the breccias exhibited end-members ranging from clast-supported textures with close-fitting, angular clasts to chaotic textures of sub-rounded components. Diagenetic sequence. The sequence of diagenetic events of palaeosols consists of the following products (Fig. 8c– d): (1) Formation of a caliche hardpan with evaporite crystals The first product observed was the precipitation of red-stained micrite. Examined samples contained polygonal components, similar to euhedral gypsum crystals

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Fig. 10. Upper Rhaetian coral limestone, Tropfbruch near Adnet (Austria). Multiphase infill of a near-surface, shallow cavity. (a) Field photograph of karstified bafflestone with scleractinian coral and coralline sponges: (4a) wackestone, (4b) bioturbated wackestone, (4c) cross-bedded packstone with erosive top, (4d) laminated wackestone with erosive top, (4e) grey wackestone, (5– 7) corroded scalenohedral and isopachous radiaxial fibrous calcite, (8) monotonous mud- to wackestone. (b) Overview of karstified microbial boundstone showing diagenetic products (5) radiaxial fibrous calcite, (6) scalenohedral calcite, (7) cement corrosion, (8) monotonous mud- to wackestone, (9a –c) fractures and 10 (stylolites). (c) Close-up view, showing radiaxial fibrous calcite (5) with scalenohedral crystals (6). Both cements were corroded (7). (d) Corroded crystals in monotonous mudstone (vadose silt).

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which replaced the lime mud prior to the lithification of the hardpan. Repeated hard pan brecciation The lithified red-stained micritic crusts were brecciated shortly after lithification at least during two events. The close fitting clasts, emphasized by abundant line contacts, indicate in-situ fracturing and/or dissolution of the hard pan without transport. Precipitation of blocky calcite Sheet cracks of the first brecciation event and moulds of roots were cemented by blocky calcite. Geopetal fill of open pores Dissolved crystals, the second generation of cracks and mould

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of roots were filled after dissolution with redstained micrite. Some sediment-filled cracks postdate cracks cemented by blocky calcite. Visible open pore space was filled after this event. The dominant pore type of the quantitatively insignificant pore system comprised non-fabric selective vugs and fractures. Fracturing Later during basin history, after the formation of sheet cracks, stratabound fractures affected the deposits of the Mahil Formation. Fracture density varies from carbonates to palaeosols (Fig. 8). The fracture patterns were not analysed in detail because they were not the main focus of this study.

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Fig. 11. Upper Rhaetian coral limestone, Tropfbruch near Adnet (Austria). Multiphase infill of shallow karst cavity. (a) Field photograph (sawed quarry wall) of karstified microbial-coral bafflestone: (4b) bioturbated wackestone, (4c–e) pack- and wackestone, (5) marine-phreatic isopachous radiaxial fibrous calcite, (8) brown, clay-rich monotonous mudstone. During the infill of sediments, Mn-minerals impregnated the roof of the cavity (yellow arrows). Note the syn-sedimentary fault which affected internal sediments (4b– e) and radiaxial fibrous calcite (5). (b) Photomicrograph of karstified microbial boundstone. Note infill of geopetal sediment (bioturbated wackestone and pack/grainstone) affected by syn-sedimentary tectonism (grey arrows), cemented fractures (9a–c) and stylolites (10). (c) Detail showing fining-upward sequences resulting in peloidal grainstone–mudstone couplets. Note thin microbial crust (yellow arrow).

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Stylolitization and creation of channel porosity The last event observed comprised low-amplitude stylolites which affected all carbonates regardless of their textures. Stylolites were locally enlarged by dissolution.

The white card technique (see methodology section above) offered the opportunity to analyse ghost structures and cross-cutting fabrics of the neomorphism at the same time (Fig. 8e–g). This simple technique facilitated the establishment of a sequence of diagenetic events for the grainstones of the cycles. Diagenetic products are as follows:

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Micritization of grains and cementation The micrite envelopes rimming the grains were observed under plane-polarized light. Carbonate cement between grains was obvious using the white card method. The morphology of the first generation cement remained enigmatic. The investigated grains (average size of 1 mm) were round or elliptical. They resembled ooids or peloids although their microstructures were destroyed. Dissolution and pore occlusion with blocky calcite The lack of ghost structures within non-skeletal or skeletal grains, and the

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Fig. 12. Upper Rhaetian coral limestone, Tropfbruch near Adnet (Austria). Multiphase infill of a deep karst cavity. (a) Field photography of cemented vug, showing (4) radiaxial fibrous calcite and scalenohedral crystals, (6) blocky calcite and (8) mudstone with corroded crystals (vadose silt). See text for details of the diagenetic sequence. Dark Spongiostromata crusts indicate that the cavity is originally a growth framework pore. (b– c) Photomicrograph of the same sequence (plane-polarized and cross-polarized light). Percentage of open porosity, indicated by blue-dyed epoxy, is low. (4) Radiaxial fibrous calcite, (5) scalenohedral calcite, (8) mudstone with corroded crystals (vadose silt). Scale in 2 mm units. (d–e) Close-up view (plane-polarized and cross-polarized light), showing (4) radiaxial fibrous calcite, (5) scalenohedral calcite, (8) mudstone with corroded crystals (vadose silt). Scale in 0.1 mm units.

(3) (4)

presence of blocky calcite suggested the dissolution of grains after cementation and the subsequent precipitation of calcite spar in the moulds. In addition, blocky calcite cemented interparticle pores to some extent. Compaction Partly collapsed components and abundant line contacts between grains indicated physical compaction of the sediments. Dolomitization Pervasive, stratiform dolomite is the most striking phenomenon of the carbonates of the Mahil Formation. The prevailing texture is unimodal non-planar dolomite (Sibley & Gregg 1987), which was most evident under plane-polarized and cross-

polarized light. Individual dolomite crystals had a ‘cloudy appearance’ and were rich in fluid inclusions. They showed no evidence of zonation under plane-polarized light. Based on observations from Jebel Akhdar, the dolomitization of the Mahil Formation has been regarded as totally fabric-destructive (e.g. Glennie et al. 1974 and references therein). Recent field work in the Saih Hatat led to the observation of textural changes of carbonates, presence of sedimentary structures and local occurrence of large fossils (Weidlich & Bernecker 2003, 2007). Using the white card method, previously undescribed petrographic

METEORIC DIAGENESIS BELOW UNCONFORMITIES

(5) (6) (7)

details were observed, including ghost structures of grains and diagenetic products. Enhancement of mouldic porosity Dissolution enhanced interparticle pores and created a fabric of well connected vugs. Fracturing See description of palaeosols Stylolitization and creation of channel porosity See description of palaeosols.

Interpretation. Micrite envelopes rimming the grains (product 1 of the diagenetic sequence of grainstones) provided evidence for early diagenetic processes in the marine-phreatic environment. If this assumption is correct at least moderate marine-phreatic cementation of the interparticle pore space might be expected for high energy grainstones. d18O values of dolomitized marine carbonates exhibit a range of 22.37 to 23.76‰ and d13C values display a range of 21.17 to 21.32‰ (see Fig. 13, Table 1). The data plotted in a narrow range which is close to the range of the isotopic composition of Middle Triassic sea water [grey field of Fig. 13; see Korte (2007)]. Considering the isotopic composition, dominance of marine facies and stratiform distribution, dolomitization may have resulted from altered marine pore waters. The interpretation as brine reflux dolomitization is less likely because positive d18O and d13C values would result from evaporation of sea water. Based on the present data, recrystallization during burial diagenesis cannot be excluded. Referring to Machel (2004) dolomite recrystallization would be regarded as insignificant if no shift of isotope data could be observed towards the field of burial dolomite. Preserved interparticle porosity was not changed by dolomitization. Mouldic porosity is believed to post-date dolomitization because no dolomite crystal faces were observed in open pores. Permeability increased through stratabound fractures (Fig. 8a–b) and channels (stylolites enlarged by dissolution, Fig. 8e). The diagenesis of the palaeosols was dominated by processes typical of a meteoric environment for the following reasons. Precipitation of the hardpan micrite (product 1 of the diagenetic sequence) took place under arid conditions with low precipitation rates. Evaporite crystals were precipitated within the capillary zone of the sediment (Warren 2006). The multiple brecciation was interpreted as being a typical feature of hard pans. Esteban & Klappa (1983) and Warren (2006) described fracturing during arid soil formation predominantly by crystallization processes. To some extent, the presence of plant rootlets might have contributed to breccia formation. Compared with marine carbonates, porosity did not increase during palaeosol diagenesis. Pore space of sheet cracks and roots has been cemented or filled with cement. Compared

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with the host rock, palaeosol formation did not improve porosity or permeability values during meteoric diagenesis. Hardly any open porosity has been found. However, thick grainstone beds may yield open pore systems which represent a mixture of interparticle and vuggy porosity. Carbon isotope values derived from the palaeosols display a range of 21.35 to 0.28‰. Two data points from the palaeosol within the cyclic sequence show positive values compared with the dolomitized marine carbonates. The depletion in 13C and the relative enrichment in 12C can be explained by the impact of isotopically light meteoric water and light soil zone CO2 (Immenhauser et al. 2002). The comparatively low d13C values of the basal Middle Triassic palaeosol which range between 20.43 to 0.28‰ plot within the range of a contemporaneous Early Triassic carbon isotopic curve of marine carbonates (21 to 5‰, see Payne et al. 2005). Neither increasing complexity of processes involved in the formation of palaeosol with time (e.g. Deutz et al. 2002), nor dolomitization models can be used to explain these values (see Warren 2000; Machel 2004 for summaries of dolomitization models). Except for one data point (basal Middle Triassic palaeosol), d18O values exhibit a range of 22.37 to 26.04‰ (see Fig. 13, Table 1). The d18O data plot in the same field as those for ooids, microbialites and cements of the Lower Triassic Rogenstein, Germany (Fig. 13). The similar negative oxygen isotope trends can be best explained with alteration/neomorphism during burial diagenesis, a phenomenon which affected many Palaeozoic and Mesozoic carbonates. These data imply that some of the blocky calcite might have been precipitated during burial diagenesis. Meteoric diagenesis enhanced the formation of a caliche hardpan and did not improve reservoir properties at the unconformity. The host rock below the unconformity did not develop a pronounced karst pore system. Although mud-dominated textures show no alteration of the pore system at all, graindominated textures have some vugs which might have resulted from meteoric diagenesis. To conclude, thick palaeosols acted as barriers rather than conduits for migrating pore fluids.

Upper Rhaetian coral limestone, Tropfbruch near Adnet, Austria Facies association. The well-known Adnet reef (Tropfbruch) has attracted the interest of palaeoecologists for many years (Scha¨fer 1979; Bernecker et al. 1999; Flu¨gel 2002). It is also famous for its marvellous changes in colour (from red to green and white), a result of repeated diagenetic processes, including occlusion of vugs by

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cements and sediment-fills. Despite their impressive appearance, investigations of diagenetic products have been rare (e.g. Schroeder & Reinhold 1997 and references therein; Weidlich et al. 2000; Reinhold & Kaufman 2009). The sawn quarry walls provide excellent exposure and diagenesis of laterally differentiated coral bioherms (stage 1), the stacked coral succession rich in debris (stage 2), and the bioclastic facies with isolated coral colonies (stage 3) can be studied. All three stages were capped by unconformities, respectively named A, B and C from bottom to top. The upper Rhaetian unconformities A and B were studied in the Tropfbruch, whereas unconformity C at the Triassic/ Jurassic boundary interval was exposed in nearby the Lienbach quarry (Figs 9–12). The three-step development documented the response of lowdiverse coral associations to low-amplitude sealevel changes, storm events and influx of platform sediment (Bernecker et al. 1999). Diagenetic sequence. Below unconformities A – C vuggy pores developed and were occluded by spatially complex generations of cements and sediments. Most complex were diagenetic products below unconformity A which were studied in detail. The spatial complexity of contemporaneous diagenetic products of carbonate rocks was described using petrogenetograms (Schroeder 1988). For this comparative analysis, the simplified description of end-members was sufficient to allow the comparison of dominant control mechanisms. In Adnet, the end-members were deep vugs (.4 m from unconformity), where cementation dominated, and near-surface vugs (distance 0–4 m), where sediment infill dominated (Fig. 9). In the following paragraph, processes prior to vug formation and after vug occlusion are described together, whereas vug occlusion processes are treated separately. Early diagenetic processes: (1) Neomorphism of metastable aragonitic skeletal grains: Aragonitic skeletal grains, like the debris of corals or molluscs, were affected by neomorphism whereas low Mg-calcite remained unaltered (Schroeder & Reinhold 1997). Neomorphism was not studied in detail. (2) Precipitation of (radiaxial) fibrous cement: Primary porosity seen in the field and in thin sections was a combination of various types, including growth framework, interparticle, intra-particle and shelter pores. Here, the first significant cement which occluded fabric selective pores is fibrous and radiaxial fibrous cement. (3) Carbonate dissolution: Formation of an irregular unconformity and creation of a connected non-fabric selective pore-system below the unconformity. Subsequent occlusion of pores varied

with distance from the unconformity. Near-surface vugs were dominated by sedimentary processes, whereas deep vugs (distance from unconformity .4 m) contained a high percentage of cement. Pore occlusion of near-surface vugs: (4) Geopetal sediment infill: The complex infill of multiple sediment generations included (4a) wackestone, (4b) bioturbated wackestone, (4c) cross-bedded packstone with erosive top, (4d) laminated wackestone with erosive top, (4e) grey wackestone. Syn-sedimentary fractures affected the sediments. (5–6) Partial pore occlusion by (radiaxial) fibrous cement and scalenohedral calcite. (7) Corrosion of radiaxial fibrous cement and/ or scalenohedral calcite. (8) Gravitational sediment infill: The sediment is crystal silt (vadose silt), monotonous mudstone or argillaceous mudstone. The sediment differed from product 4 with respect to texture and sedimentary structures. Pore occlusion of deep vugs comprised: (4) Partial occlusion by radiaxial fibrous calcite (5)–(6) Partial occlusion by scalenohedral calcite and blocky calcite (7) Corrosion of calcite cements (8) Infill of monotonous, partly argillaceous mudstone, sometimes with crystal silt After the complete pore occlusion, late diagenetic products include: (9) Fracturing: Three generations of fractures were observed. Fractures were cemented by calcite (Fig. 10b). (10) Stylolitization: The latest observed phenomenon was the formation of low amplitude stylolites (Fig. 10a). Interpretation. The dissolution and cementation processes described here caused the most complex pore system of the three studies with respect to size and pore occlusion. The pore classes ranged from micro- to macroporosity (classification of Luo & Machel 1995). Based on field data and petrographic observation, meteoric karst diagenesis was regarded as the ultimate control mechanism of nonfabric selective pores (Schroeder & Reinhold 1997; Bernecker et al. 1999; Weidlich et al. 2000). The unconformities were interpreted as resulting from karst events because of: the significant relief of unconformities; rillenkarst features at top of the unconformities; abundant vugs truncating the primary fabric; and the abundance of dissolved aragonitic skeletal grains. The observed features correspond with the criteria listed for karst surfaces by Hillga¨rtner (1998). Karst phenomena affected other Triassic reefs in the Northern Calcareous Alps (e.g. Satterley et al. 1994) and formed an important feature of the Lofer cycles.

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Fig. 13. Stable isotope data (grey areas represent possible composition of sea water, see text for references). (a) Lower Triassic Rogenstein, Calvo¨rde Formation, Buntsandstein Group. Locality Harliberg, Germany. (b) Middle Triassic Mahil Formation, Locality Wadi Aday, Saih Hatat, Sultanate of Oman. (c) Upper Rhaetian coral limestone, Locality Tropfbruch near Adnet, Austria.

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Table 1. Table showing stable oxygen and carbon isotope data

d13C PDB

d18O PDB

Material

Lithostratigraphy

1A-1 3 4-1 7-1 8-1 1A-2 5-1 5-2 7-2 8-2 BB01-01 BB01-02 BB01-03 BB01-03 D02ost1-1 D02ost1-2 D02ost1-3 D02Ho1-1 D02He1-01 D02He1-02 D02He1-03 D02He1-04 Be-aP-01-01 Be-aP-01-01 Be-aP-01-02 1A-3

20.75 22.20 20.63 20.30 21.24 22.10 0.25 0.07 0.26 20.10 0.77 0.72 0.74 0.71 20.29 20.25 20.19 21.44 20.74 20.71 20.40 20.35 0.64 0.60 0.59 0.25

26.03 25.35 26.41 26.34 6.38 25.13 25.13 25.73 25.79 25.75 26.58 26.54 26.25 26.28 26.65 26.77 26.79 26.23 26.98 26.54 26.44 26.53 26.75 26.63 26.76 25.52

blocky & poikilotopic calcite blocky & poikilotopic calcite blocky & poikilotopic calcite blocky & poikilotopic calcite blocky & poikilotopic calcite fine-grained layer ooid ooid ooid ooid ooid ooid ooid ooid ooid ooid ooid ooid ooid ooid ooid ooid ooid ooid ooid top microbialite

L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group L. Buntsandstein Group

Chronostratigraphy Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic Early Triassic

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Sample

20.50 24.70

25.72 25.98

microbialite microbialite

L. Buntsandstein Group L. Buntsandstein Group

Early Triassic Early Triassic

65 66 67 69 68 70 71 OW49new OW45new 72 73 74

21.20 21.32 21.17 21.17 21.35 21.16 20.98 0.25 0.28 20.43 21.66 21.55

22.37 23.33 22.84 23.76 23.53 24.63 24.12 3.65 24.91 24.73 26.04 24.76

wackestone mudstone mudstone packstone palaeosol palaeosol palaeosol basal MT palaeosol basal MT palaeosol basal MT palaeosol black calcite black calcite

Mahil Formation Mahil Formation Mahil Formation Mahil Formation Mahil Formation Mahil Formation Mahil Formation Mahil Formation Mahil Formation Mahil Formation Mahil Formation Mahil Formation

Middle Triassic Middle Triassic Middle Triassic Middle Triassic Middle Triassic Middle Triassic Middle Triassic Middle Triassic Middle Triassic Middle Triassic Early Triassic Early Triassic

NKW8A-1 NKW8A-2 K2-1 NKW8A-3 NKW8A-4 NKW8A-5 NKW8A-6 NKW8A-7 NKW8A-8 K2-2 K2-3

3.05 3.15 2.57 3.17 3.06 2.79 2.15 2.44 20.35 2.24 2.32

20.47 20.77 21.24 20.72 22.23 27.98 27.88 27.65 26.81 21.30 21.03

reef reef reef marine-phreatic cement marine-phreatic cement blocky calcite blocky calcite eroded blocky calcite matrix sediment sediment

U. Rhaetian Reef U. Rhaetian Reef U. Rhaetian Reef U. Rhaetian Reef U. Rhaetian Reef U. Rhaetian Reef U. Rhaetian Reef U. Rhaetian Reef U. Rhaetian Reef U. Rhaetian Reef U. Rhaetian Reef

Late Triassic Late Triassic Late Triassic Late Triassic Late Triassic Late Triassic Late Triassic Late Triassic Late Triassic Late Triassic Late Triassic

METEORIC DIAGENESIS BELOW UNCONFORMITIES

D02Ho2-01 D02Ho2-03

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The stable isotope data from these sediments support the petrographic observations and suggest that meteoric diagenesis was the driving mechanism was for the formation of unconformities and the vugs (Fig. 13, Table 1). Reef sediments (d13C ¼ 2.57 to 3.15‰; d18O ¼ 20.47 to 21.24‰), radial fibrous calcite (product 2 of nearsurface vugs, d13C ¼ 3.06 to 3.17‰; d18O ¼ 2 0.72 to 22.23‰) and geopetal sediment infill (product 4 of near-surface vugs, d13C ¼ 2.24 to 2.32‰; d18O ¼ 21.03 to 21.30‰) plotted within the field of probable isotopic composition of Upper Triassic sea water. The isotopic composition of vug-filling crystal silt, monotonous mudstone and argillaceous mudstone (product 8 of near-surface vug, d13C ¼ 20.35‰; d18O ¼ 26.81‰) showed a negative carbon isotope value compared with the other samples. This evidence pointed towards an increase of the light isotope 12C during the formation of soils and, thus, meteoric conditions during sedimentation. Finally, blocky calcite (product 5 of deep vugs, d13C ¼ 2.15 to 2.70‰; d18O ¼ 27.65 to 27.98‰) was characterized by negative oxygen values. This trend resembled product 8 and was interpreted as influence of burial diagenesis. The well connected pore systems below the unconformity were completely filled. Shallow cavities were filled with multiple generations of internal sediments, whereas deep karst cavities were cemented. This pattern indicated good connectivity of the karst pores in the upper part of the system. The bioturbated wackestone of the shallowest cavities indicated a marine environment with enough oxygen for coelobites (organisms living in cryptic environments). Cross-bedded packstone with erosive top and laminated wackestone pointed to varying hydrodynamic conditions in the karst system. Isopachous rims of marine-phreatic calcite were relatively thin in shallow cavities. Conversely, deep cavities were almost completely filled with a characteristic sequence of cements. Corrosion of various cements in deep and shallow cavities is considered to have been a contemporaneous process, which may have occurred during the next karst event (unconformity B), when carbonateundersaturated fresh water percolated through the karst system during sea-level lowstand. Topography of unconformity A and depth of epikarst suggested a sea-level fall of up to 15 m (Bernecker et al. 1999; Fig. 9). At unconformity B, a fall of sea-level of at least 3–4 m has been estimated based on the karst topography of the unconformity. The penetration depth of carbonate under-saturated meteoric pore water probably exceeded 17 m while the karst system of unconformity A was reactivated again. The depth of epikarst at unconformity C was probably less than 2 m.

In the massive reefs, the fractures cannot be classified as stratabound or non-stratabound. Fractures affected reef rock and karst cavities in a similar way. All fractures were either cemented or filled with internal sediment and they have been cut by two generations of late-diagenetic stylolites.

Interpretation Once the mechanisms and nature of the processes contributing to the generation and occlusion of porosity at and below these described unconformities had been identified, the analysis proceeded in the following manner: (1) transfer of the presented sedimentological and diagenetic data into petrophysical data; (2) comparison of the results with petrophysical classes (Lucia 1995); and (3) finally the interpretation of the control factors on meteoric diagenesis.

Data transformation into petrophysical classes The open porosity of representative samples was quantified using image analysis from thin sections. These thin sections were impregnated with bluedyed epoxy and permeability was calculated using the equation of Mowers & Budd (1996). This was an important step because the transformation of geological data to rock-fabric characteristics is focused on porosity and permeability (Lucia 1995). In terms of petrophysical needs, classification of pore space had to be simplified. As a consequence, porosity was treated either as interparticle or vuggy. Interparticle pore space was subdivided into petrophysical classes sensu Lucia (Fig. 14), where class 1 comprised grainstone textures with excellent pore connectivity (grain size varies between 500–100 mm), class 2 packstone textures with good pore connectivity (grain size varies between 100–20 mm) and class 3 pack-, wacke- and mudstone textures with poor pore connectivity (grain size ,20 mm, Lucia 1995). Interparticle, intercrystalline and growth framework porosity of Choquette & Pray (1970) were all treated as interparticle porosity. Fabric selective and non-fabric selective pore types of Choquette & Pray (1970), including intra-particle, mouldic, shelter, fenestral, cavern, breccia, fracture and channel porosity, were all considered as vuggy porosity in this classification. Vuggy porosity was further subdivided into separate-vug pores and touching-vug pores (Lucia 1995). Permeability decreased from touching-vug porosity to separatevug porosity. The transfer of data into petrophysical groups provided the basis to interpret the porosity and permeability properties of carbonates below

METEORIC DIAGENESIS BELOW UNCONFORMITIES

311

Fig. 14. Porosity and permeability cross plot from image analysis data of case study samples compared to petrophysical classes of Lucia (1995).

unconformities (Fig. 14). The case studies allowed the differentiation of three scenarios. (1)

(2)

Increase of porosity and permeability and their preservation through time: Open porosity and permeability data of the Middle Triassic oolites of the Mahil Formation (locality 2) and of the Early Triassic Rogenstein (Calvo¨rde Formation, Buntsandstein Group, Germany, locality 1) plotted close to petrophysical class 1 or fall within class 1. This was the result of the presence of mixed pore systems consisting of interparticle and touching-vug pores. The permeability of the pore systems of the oolites further increased because of meteoric dissolution. The stacked cycles of both formations with repetitive succession of oolites allow the assumption that facies associations were laterally persistent. It can be concluded that meteoric diagenesis compensated the mediocre porosities and improved reservoir properties and provided conduits for formation waters during basin history. Increase of porosity and permeability and subsequent pore system occlusion: The upper Rhaetian coral limestone (Tropfbruch near Adnet, Austria, locality 3) is a paradigm of multiple karst events and significant carbonate dissolution. Touching-vug pore space dominated the system and was the result of growth framework pores enlarged by dissolution, mouldic pores and vugs. Interparticle porosity was only of local importance. The excellent communication of the touching-vug pore

(3)

space gave rise to its subsequent occlusion with cement and sediment. The open porosity and permeability data obtained from large thin sections (size is 5 cm  7.5 cm; core plug data did not exist) consequently plotted in the lower left corner of the diagram presented in Figure 14. It can be concluded that the karst system changed from a conduit to a barrier slowing down the movement of formation waters. Decrease of porosity and permeability and creation of a barrier for pore fluids: Meteoric diagenesis responsible for the formation of the palaeosols did not improve porosity and permeability of the Middle Triassic at the unconformities of the Mahil Formation (Akhdar Group, Sultanate of Oman, locality 2). Pedogenic processes triggered the precipitation of tight caliche carbonates and clay minerals were enriched in the lower part as a result of host rock dissolution. Transformation of data into petrophysical classes in the lower left corner is a visualization of the poor reservoir properties (Fig. 14) and the decrease of porosity and permeability.

Control mechanisms of meteoric diagenesis The interpretation of the mechanisms controlling the dimensions of karst pore systems and their preservation is challenging. Key factors among many others are the time span of sub-aerial exposure, the amplitude of sea-level change, the position of palaeo-water table and precipitation rates (climatic variation).

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During the last 15 years, orbital forcing has been used in the literature to explain the nature of cyclic carbonate successions and to constrain the time spans involved in the formation of their unconformities. It has been proposed that 100 ka eccentricity cycles prevailed during ice-house climate while 20 ka precession cycles were more important during times of greenhouse climate, such as during the Triassic (Read et al. 1995). Among the various proposed models explaining controls of the lithologically obvious cycles of the Buntsandstein Group (locality 1), orbital forcing, under consideration of precession cycles, has been suggested (see Geluk & Ro¨hling 1997 and further references therein). The time period represented by the shallowing-upward cycles of the Mahil Formation has not yet been investigated. However, it has been proposed that similar cycles of the contemporaneous Latemar platform were triggered by Milankovitch cycles (e.g. Goldhammer et al. 1990) as well, notably 20 ka precession cycles with superimposed 100 and 400 ka eccentricity cycles (Read et al. 1995). Karst events of the Upper Triassic Wilde Kirche reef, a buildup similar to the Adnet patch reef, have been also explained by orbital forcing, notably precession cycles (Satterley et al. 1994). It can be concluded that geologically similar time spans, of significantly less than 20 ka years, were probably involved in the formation of Triassic karst unconformities of the presented case studies. The amplitude of sea-level changes can be expected to be low during greenhouse climate due to the absence of polar ice caps which store large amounts of water during lowstands of sea-level (Read et al. 1995). The distribution and abundance of climate-sensitive evaporites and tropical carbonates pointed towards greenhouse conditions during the Triassic (Frakes et al. 1992). This interpretation has been supported by modelled atmospheric CO2 (e.g. Berner 1991), indicating an increase of atmospheric carbon dioxide in the atmosphere and a warm climate throughout the Triassic. The mechanisms of sea-level change during a greenhouse climate were consequently subject to controversial discussions. Among several proposed solutions, storage of water in aquifers or lakes has been considered (see Preto et al. 2001 for further references). Whatever the mechanism, the amplitude of sea-level changes was similar for the three case studies and did not cause the variations of karst diagenesis of the case studies. Fluctuations of the palaeo-water table may have been a mechanism. Reliable data which allowed the interpretation of this factor exists only for locality 3. The interpretation was not possible with the available data set. As mentioned earlier, during the Triassic there was a greenhouse climate. Assuming uniform

climatic conditions is an oversimplification. Terrestrial environments of the supercontinent Pangaea are believed to have undergone extreme seasonal temperature variations (Crowley 1994). This may have led to an underestimated control on early diagenesis of land-locked basins and marginal carbonate platforms. The size of the Pangaean supercontinent may have caused especially intense climate events such as so-called mega-monsoons (Kutzbach & Gallimore 1989) which led to drastic variations in the patterns of rainfall density. These intense climate conditions, with monsoon events, may have led to the intensification of meteoric dissolution. Modelled palaeoclimate data (Golonka et al. 1994) assisted in the interpretation of regional heterogeneities of rainfall density and the intensity of meteoric diagenesis during Triassic times (Fig. 15). Variations of rainfall density were plotted as winter and summer maps of the northern hemisphere. Areas of over 70% probability of humid climatic conditions were marked in yellow in Figure 15. If the locations 1 to 3 of this study

Fig. 15. Palaeogeographic and palaeoclimatic maps. (a) Lower–Middle Triassic palaeogeography and palaeoclimatology. The distribution of monsoon areas was modelled for the winter northern hemisphere. All maps modified after Golonka et al. (1994). (b) Upper Triassic palaeogeography and palaeoclimatology. The distribution of monsoon areas was modelled for the summer northern hemisphere. Key for palaeogeography and palaeoclimatology colour codes: dark blue, open ocean; light blue, shallow shelf; green land; yellow areas with over 70% probability of favourable conditions for equatorial rainfalls (monsoon).

METEORIC DIAGENESIS BELOW UNCONFORMITIES

are plotted on these maps, it becomes evident that localities 1 (Rogenstein, Germanic Basin) and 2 (Mahil Formation Oman) have a distance of 250 –500 km to modelled monsoon systems. Conversely, locality 3 (Upper Triassic Adnet reef ) was touched by a monsoon system. As a consequence precipitation rates would be expected to have been highest in the vicinity of locality 3 and lowest for localities 1 and 2. A positive correlation between temporarily high precipitation and size of karst system became evident based on this comparison. The observed differences of meteoric dissolution could be explained in an unconventional way as driven by monsoon cells in the following way: (1) The combination of modelled palaeoclimate data with the observation of a big karst system (Adnet, locality 3) gave rise to the interpretation that the study area was situated close to a monsoon system. As a result of the large size, the karst system was occluded during basin evolution. (2) The combination of modelled palaeoclimate data with the observation of a small karst system (oolites of the Rogenstein, Calvo¨rde Formation, locality 1 and oolites of the Mahil Formation, locality 2) suggested that the study areas were not touched by monsoons. The limited size of the pore systems of the oolites of the Calvo¨rde Formation (case study 1) and distance to a monsoon system (Fig. 15) led to the assumption that rainfall provided enough freshwater during sub-aerial exposure to enhance the size of the existing pore system of oolites. It was concluded that dissolution was not significant enough to create a well-connected vuggy karst system. The smaller size of the pore system and narrow pore throat diameters was one factor responsible for the partial occlusion of pores. (3) The combination of modelled palaeoclimate data with the observation of tight caliche hard pans of the Mahil Formation (case study 2) suggested that local precipitation was moderate and affected by monsoon systems. The carbonates of the palaeosols indicated that evaporation exceeded precipitation and caliche-like carbonates formed in the capillary zone. Palaeosols of the Mahil Formation formed therefore local barriers for pore fluids in the study area.

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preservation through time; the increase of porosity and permeability and subsequent pore system occlusion; and the decrease of porosity and permeability and creation of a barrier for pore fluids. The observed differences in the size of the pore systems and the presence of various diagenetic products generated within this pore system could not be interpreted in a conclusive way by variations in sea-level changes or amplitude changes of sealevel. The combination of field data, petrographic observation, stable carbon and oxygen isotope data and porosity and permeability data enabled the characterization of porosity and permeability and its conversion into petrophysical classes. Comparison of these data with palaeogeographic maps and modelled palaeoclimate simulations highlighted the importance of rainfall patterns in meteoric dissolution. It is suggested that large karst systems have been generated in areas where monsoon systems were present during the Triassic. However, large karst systems were concealed by subsequent sedimentation and cementation processes. Conversely, smaller karst systems were formed in areas where semi-arid climatic conditions prevailed without the impact of monsoon events. These karst systems were more likely to remain open during basin evolution. In an arid climate, if meteoric diagenesis was dominated by the formation tight caliche hardpans at unconformities, these will have acted as barriers for pore fluid migration. This comparative analysis has shown that the data from diagenetically-altered carbonate formations need to be integrated with the full suite of sedimentological, structural, palaeoclimatic and tectonostratigraphic studies to improve our understanding of the controls on rock fabrics and reservoir properties that we observe today. This study was partly financed by the German Science Foundation (project We1804/8-1,2). I thank D. Lowry (Royal Holloway, University of London, United Kingdom) and H. Erlenkeuser (Leibniz Laboratory, Kiel University, Germany) for stable isotope measurements. M Bernecker (Universita¨t Erlangen-Nu¨rnberg) improved with valuable comments the manuscript. The help and enthusiasm of the editors of this volume, F. S. P. van Buchem, K. D. Gerdes and M. Esteban, is greatly acknowledged. Constructive reviews of A. E. Csoma, B. Vincent and K. D. Gerdes significantly strengthened the manuscript. Special thanks go to the Ministry of Commerce and Industry in the Sultanate of Oman, especially the Director General of Minerals, Hilal bin Mohamed Al-Azri for administrative support.

Conclusions Focusing on three case studies, diagenetic processes of carbonates below Triassic unconformities were studied. Three different scenarios were identified: the increase of porosity and permeability and their

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Porosity development, diagenesis and basin modelling of a Lower Cretaceous (Albian) carbonate platform from northern Spain I. ROSALES1* & A. PE´REZ-GARCI´A2 1

IGME-Instituto Geolo´gico y Minero de Espan˜a, La Calera 1, 28760 Tres Cantos, Madrid, Spain

2

REPSOL-YPF, Exploracio´n y Produccio´n, Paseo de la Castellana 280, 28046 Madrid, Spain *Corresponding author (e-mail: [email protected]) Abstract: This paper describes an integrated field, petrographic and geochemical study of an Albian carbonate platform from the Basque-Cantabrian basin (northern Spain). It examines the distribution and evolution of porosity and the relationships between facies variations, sequences and variations in diagenesis. These platform carbonates were deposited on the footwall crests of active tilted blocks formed during continental rifting related to the Mesozoic opening of the Bay of Biscay and North Atlantic. The studied units document an early Albian aggradational steep-sloping platform and a late Albian, low-gradient expansive shelf separated by a hiatal unconformity spanning the middle– early late Albian (c. 5 Ma). The late Albian platform unit also exhibits a major internal unconformity. Platform geometries and facies architecture were mainly controlled by tectonics, hydrodynamic energy level and water depth. Petrographic, cathodoluminescence and geochemical analyses suggest that early meteoric diagenesis developed during sub-aerial exposure in strata below these two major unconformities. The platform carbonates have been affected during burial by a number of diagenetic processes that include four phases of dissolution, several fracture generations, and six cement sequences with development of at least 13 calcite and dolomite cement zones (Z0– Z12). A contrasting diagenetic response from the different platform environments illustrates the role of primary sediment composition and unconformity development in controlling porosity and cement distribution. Limestone stabilization and cementation were relatively early processes that were mostly completed within the first kilometre of burial depth beneath the depositional surface. Below this burial depth, fluid circulation was concentrated along restricted pathways (fractures and fault zones). Migration of hydrothermal-related fluids along fault zones created localized dolomite patches and large-scale porosity associated with cavities and collapse breccias, but did not significantly increase the small-scale porosity.

The Aptian –Albian shallow-water carbonate platforms (Urgonian platforms) of the Basque– Cantabrian basin (BCB) of northern Spain (Fig. 1) host one of the largest metal ore bodies of Europe, and although they do not host known hydrocarbon fields, they exhibit petroleum shows both at the surface (Rosales et al. 1992) and in the subsurface (e.g. Mar Canta´brico M wells offshore Asturias). These platform carbonates have been intensively studied, particularly with respect to their general stratigraphy and sequence stratigraphic organization (e.g. Rat 1959; Pascal 1984; Garcı´a-Monde´jar 1990; Ferna´ndez-Mendiola et al. 1993; Garcı´a-Monde´jar & Ferna´ndez-Mendiola 1993; Rosales et al. 1994; Neuweiler 1995; Go´mez-Pe´rez et al. 1999; Rosales 1999) as well as for the metalogenetic aspects (e.g. Velasco et al. 1994, 1996, 2003; Bustillo & Ordon˜ez 1995; Simon et al. 1999). Despite these studies, relatively little attention has been directed toward the diagenesis and porosity evolution of these platform carbonates and their potential as

carbonate reservoirs. The aims of this study were the following: (1) to delineate the diagenetic history of the Albian shallow-water limestones of the Ramales area, which is one of the largest carbonate-hosted metal ore provinces in the BCB; (2) to determine the main controls on the distribution of diagenetic products and their effects on reservoir properties; and (3) to relate these diagenetic processes to the depositional, thermal and burial history of the area. The stratigraphy and the depositional and subsidence histories of the Ramales platform have been discussed in detail by Lo´pez-Horgue (2000) and Garcı´a-Monde´jar et al. (2005). Our study is limited to the late early Albian to late Albian succession of these platform carbonates. This succession is of interest to petroleum companies due to possible development of secondary porosity associated with sub-aerial exposure and local dolomitization. This study presents new diagenetic data from these Lower Cretaceous platform carbonates and uses these data

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 317–342. DOI: 10.1144/SP329.13 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Schematic palaeotectonic map of the north Spanish margin and Bay of Biscay during the Albian showing the location of the Basque-Cantabrian basin and the distribution of the main carbonate platform areas (modified after Garcı´a-Monde´jar & Ferna´ndez-Mendiola 1993).

to assess the hydrocarbon reservoir potential of these sequences. Finally, the diagenetic evolution developed for the Ramales carbonate platforms is integrated with the established tectono-stratigraphic framework for the area to explain the spatial and temporal distribution of porosity in the platforms. The integration of these results with the regional subsidence and thermal history allows this diagenetic study to be set in the context of the basin evolution.

Regional setting The Basque–Cantabrian basin is a Mesozoic– early Cenozoic extensional basin that was uplifted and folded in the Palaeogene during the Alpine orogeny. The geodynamic evolution of the BCB was related to the kinematics between the European and Iberian plates, which are closely linked to the extensional processes during the opening of the North Atlantic Ocean and the Bay of Biscay. The basin records three major tectonic events: (1) PermoTriassic rifting; (2) rifting and opening of the Bay of Biscay during the late Jurassic –Early Cretaceous; and (3) the Pyrenean orogeny during the Eocene –Oligocene. Therefore, the present structural framework of the BCB reflects the overprinting of the Alpine compression onto the Mesozoic extension and strike-slip phases. The inverted basin constitutes the western prolongation of the Pyrenean folded belt and extends along onshore and offshore areas of central-northern Spain (Fig. 1). The basin records a thick pile of sediments that measure more than 10 km in the most subsiding areas. In the study area, the outcropped Mesozoic succession records about 4000 m of continental to shallow marine Triassic to Cenomanian deposits. The original

thickness of the Late Cretaceous and Palaeogene sediments in the study area is unknown because of post-orogenic erosion. In the BCB, a phase of strong tectonic subsidence existed during Albian times, as tension caused by intraplate stress created differential fault-driven subsidence and active rift tectonics (Garcı´aMonde´jar et al. 1996). In this tectonically active setting, thick masses of shallow subtidal limestones (the so-called Urgonian limestones; Rat 1959) were deposited on the footwall crests of tilted fault blocks. Shallow-water carbonates passed abruptly into deeper-water marls and mudstones deposited in hanging-wall depocentres. One of the largest Albian carbonate platforms of the BCB was the Ramales platform in the Cantabrian region (Fig. 2), with a complex sequential and diagenetic evolution. Commonly, the Urgonian platforms of the central BCB evolved with time into four major phases of carbonate platform development: carbonate ramp (early Aptian),distally-steepenedramp(late Aptian – earliest Albian), progradational-aggradational steep-sloping platform (early Albian), and carbonate bank (late Albian). These stacked major platform intervals are vertically separated by regional unconformities that involve platform exposure and/or drowning usually accompanied by siliciclastic influx (Garcı´a-Monde´jar 1990; Ferna´ndezMendiola et al. 1993; Rosales 1999; Go´mez-Pe´rez et al. 1999). Each platform phase may be subdivided into several minor depositional sequences. Platform geometries and internal facies architecture within each platform phase were mainly controlled by tectonics, relative sea-level changes, and hydrodynamic and climatic factors (e.g. Pascal 1984; Garcı´aMonde´jar 1990; Neuweiler 1993; Rosales et al. 1994; Go´mez-Pe´rez et al. 1999; Rosales 1999). Palaeoclimatic indicators in the area suggest warm and humid conditions during Aptian and Albian times, as supported by the lack of evaporitic deposits, common storm deposits and frequent fossil wood and amber fragments in coeval deltaic facies. Local syn-sedimentary tectonics and relative sea-level changes controlled both the geometry of the platform-to-basin transition and the accommodation, which in turn controlled facies distribution. This paper focuses on the depositional and diagenetic evolution of the two younger carbonate phases (early Albian and late Albian) developed on the southeastern margin of the Ramales platform in the environs of the type locality at Ramales de la Victoria (Fig. 3).

Methodology Four correlative stratigraphic sections (Fig. 3a, b) 240–450 m thick were logged and sampled at a

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Fig. 2. Palaeogeographic reconstruction of the Basque-Cantabrian basin during the early Albian, with palinspastic restoration from surface and subsurface data (modified from Garcı´a-Monde´jar 1990).

metre-scale along a c. 10 km long platform-to-basin transition. Sixty-three standard-size, uncovered, polished thin sections were prepared and partially stained with a mixed solution of Alizarin Red-S and potassium ferricyanide (Dickson 1966) to aid in the identification of ferroan and non-ferroan phases of calcites and dolomites. Thin sections were examined under petrographic and cathodoluminescence (CL) microscopes for identification of mineralogy, diagenetic textures and mineral paragenesis. Selected thin sections were point counted (c. 300 counts per sample) for modal analysis in order to evaluate quantitatively the mineral composition and porosity evolution (creation and destruction of porosity). The methods used in this study are only capable of evaluating porosity at thin-section scale (,1 cm). CL analysis, in conjunction with stable isotope analysis, has provided a history of porosity occlusion. The CL study was carried out on a Nuclude ELM-3R luminoscope under 10 –15 kV beam potential, 0.5 mA beam current, and 0.05–0.1 torr pressure. After CL evaluation, different generations of calcite and dolomite cements were analysed for their d13C and d18O composition. Twenty-two samples for stable isotope analyses were powdered from polished rock chips (minimum c. 1 mg) using

a microscope-mounted dental drill supplied with tungsten bits. Stable-isotope data were collected from the CO2 gas liberated by the carbonate samples after reacting with 100% phosphoric acid, using a VG SIRA-9 mass spectrometer at the University of Salamanca (Spain). Data are expressed in ‰ relative to the Pee Dee Belemnite (PDB) standard. The analytical precision based on replicate analyses of a standard is better than 0.2‰ for d13C and d18O. One dimensional basin modelling using Platte River’s BasinMod software was used to establish the burial and thermal history of the study area. Data used to calibrate the thermal model were obtained from vitrinite reflectance and visual kerogen techniques. Measurements were made on carefully polished blocks of either kerogen concentrates or whole rocks by the DGSI laboratories.

Stratigraphic framework The Aptian to Albian succession at Ramales records nearly 900 m of shallow-water carbonates that are correlatable to the south and east with marls, marly limestones and siliciclastics deposited in slope, basinal and siliciclastic shelf environments (Fig. 3c, d). The general stratigraphic framework,

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Fig. 3. (a) Aerial photograph showing the study area and location of the four studied stratigraphic sections. Aerial images sourced from the Cantabria Government. (b) Simplified surface geological map of the Albian deposits of the Ramales area. The red square with dashed line marks the study area shown in (a), and the numbered red thick lines indicate the location of the four measured sections. 1, Ojebar section; 2, Ranero section; 3, Cueto Misario section; 4, Fuente Frı´a section. (c) Chronostratigraphic chart for the main lithostratigraphic units (ages of the lithostratigraphic units based on R.W. Scott, pers. comm.; Lo´pez-Horgue et al. 1994, 2000; Aranburu 1998; Garcı´a-Monde´jar et al. 2005). (d) Stratigraphic cross-section based on mapping and logged sections of the area shown in (a) and (b).

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and main depositional systems and sequences, have been presented previously (Go´mez 1989; Lo´pezHorgue et al. 1994; Lo´pez-Horgue 2000; Garcı´aMonde´jar et al. 2005). Figure 3(c and d) shows a simplified stratigraphy of the lower to upper Albian succession in the study area, which has been adapted to follow the most recent terminology by Garcı´a-Monde´jar et al. (2005) for the main lithological units. Benthic foraminifers (R. W. Scott pers. comm.; Lo´pez-Horgue et al. 1994, 2000) and ammonites from the basinal units (Aranburu 1998) provide the chronostratigraphic framework for this study. Two major carbonate depositional systems developed at Ramales during the Albian (Fig. 3d). A lower Albian aggradational steep-sloping carbonate platform (Ranero limestones, c. 450 m thick, Fig. 3d) and an upper Albian low-gradient widespread carbonate bank (Sopen˜a limestones 120 –240 m thick, Fig. 3d). They are separated by a regional unconformity that, according to available biostratigraphic data based on benthic foraminifers (orbitolinids), spans the middle Albian to early late Albian (c. 5 Ma, Fig. 3c; Garcı´a-Monde´jar et al. 2005). Basinward equivalents to these platform units are basinal marls, downslope carbonate breccias and basin infilling siliciclastic deposits (Fig. 3d). The Ranero limestones grade basinward into the Cuadro marls (Garcı´a-Monde´jar et al. 2005), a c. 300 –500 m thick unit of basinal marls, marly limestones and downslope carbonate breccias. The Cuadro marls are overlain by the siliciclastic basin-infilling units of the Rı´o Calera and Valmaseda Fm (Fig. 3d). The Rı´o Calera unit (early late Albian) is made of marls, lutites and siltstones with intercalated sandstone turbidites and carbonate megabreccias deposited on a siliciclastic slope. The Valmaseda Fm (late Albian –early Cenomanian) is a shelfal siliciclastic unit deposited in storm-dominated shoreface to offshore environments (Pe´rez-Garcı´a et al. 1997). It can be informally subdivided into three subunits: (1) a lower siliciclastic subunit that pinches out toward the Ranero platform margin (Fig. 3d); (2) a middle siliciclastic subunit that grades and interfingers laterally with carbonate strata of the Sopen˜a limestones; and (3) an upper siliciclastic unit that covered the entire carbonate platform. The Rı´o Calera unit and lower Valmaseda subunit progressively thin toward the Ranero platform margin and thicken downdip indicating a rapid infilling of the basin (Fig. 3d).

Depositional facies and sequences The two studied platform units are very well exposed in the study area, which allows the definition of depositional sequences and facies distribution to

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be made along continuous shelf-to-basin outcrop exposures (Fig. 3).

The Ranero limestones Sedimentary facies of the Ranero limestones have been grouped into platform interior, platform margin and platform slope facies assemblages on the basis of rock textures, skeletal components, fossil content, bedding and geometric relationships. The platform interior environments (Ojebar section, Fig. 4) are characterized by micritic limestones with rudist assemblages. Lithofacies include rudistmiliolid wackestones and floatstones with intercalated rudist and Chondrodonta biostromes, peloidal-oncoidal wackestone-packstones, and organic buildups up to 9 m high. These lithofacies were deposited in a low-energy, muddy, restricted lagoon (Lo´pez-Horgue 2000). Towards the platform margin, lagoonal lithofacies (Ranero section, Fig. 5) grade from structureless to cross-bedded skeletalorbitolinid packstones and grainstones (Fig. 6a, b), coral-skeletal grainstones (Fig. 6c, d), and fossiliferous patch reefs with rudists, corals, nerineids, orbitolinids and braquiopods. These deposits represent a higher-energy, shallow shelf-margin environment. Downslope, these marginal facies grade into large-scale microbial-sponge-coral-algae mud mounds, up to 70 m high, developed on the platform slope below wave-base (Neuweiler 1995). These micritic buildups proliferated during this time in platform slope and margin environments of many platform areas of the BCB (Garcı´a-Monde´jar & Ferna´ndez-Mendiola 1995; Neuweiler et al. 1999), and they usually have algal (Bacinella irregularis) or microbialite/sponge frames (Neuweiler 1993; Rosales et al. 1995). The accumulation and preservation of the mud mounds on the slope has been related to the high rates of tectonic subsidence that occurred during the early Albian in this part of the basin (Garcı´a-Monde´jar et al. 1996, 2005).

The Sopen˜a limestones The sedimentary facies of the Sopen˜a limestones (Figs 4, 5, 7 & 8) have been grouped in this study into inner shelf, wave/storm-dominated sand shoal and outer shelf facies based on rock textures, sedimentary structures, fossil content, mapping of lithofacies and physical correlation in the dip direction (Fig. 9). Inner shelf facies consist of two facies associations: skeletal and foraminiferal, poorlysorted, massive to poorly laminated packstones and grainstones, and coral-rudist bioherms (Fig. 8a). The shallow-water biota and the poorly-developed sedimentary structures are interpreted as shallow subtidal sands and patch reefs developed in bankprotected areas which were episodically agitated.

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Fig. 4. Logged stratigraphic section in Ojebar showing lithofacies, distribution of age-diagnostic benthic foraminifers, location of the analysed samples and the stratigraphic distribution of CL cement zones related to dissolution events. The presence of a particular cement zone in primary or secondary porosity for each sample is represented by a black dot.

The wave/storm-dominated sand shoal facies complex comprises three different facies associations. (1) Fine sand-sized, peloidal-foraminiferal packstones and grainstones with oscillatory ripples, wavy stratification and millimetre-scale marly and clay laminae (Fig. 8b, c). Red algae, chaetetids, orbitolinids and peloids played an important role in the production of these carbonate sands. Thickshelled oysters and platy corals (microselenids) frequently colonized the rippled calcarenite tops within the marly drapes. (2) Medium to coarse-grained grainstones with wavy muddy laminae are locally, massive, tabular and may be colonized by colonial corals and oysters. (3) Coarse sand-sized, cross-bedded, skeletal-intraclast-oolitic grainstones (Fig. 8d, e) occur with cross to hummocky cross stratification, flat bases and rippled tops. These latter beds occur in continuous sheet-like to lensoid bodies several hundred metres long and 3–6 m thick. These three lithofacies stack in fining-upward successions that are 3–10 m thick.

Typically, their lower contacts are sharp and erosional, and may be overlain by carbonate rudstones and gravels containing sparse, unoriented quartzite pebbles up to 10 cm long (Fig. 8f ). They are overlain by the cross-bedded, coarse-grained carbonate sands, and finally by fine-grained crosslaminated calcarenites with the clay caps colonized by oysters and platy corals. In some sequences, wave-rippled calcarenites directly overlie the erosional surface. The presence in these deposits of erosive surfaces and hummocky cross-stratification suggests that the environment was periodically affected by storm currents. Current energy may have intensified during seasonal storms and caused erosional surfaces, the redistribution of carbonate gravels and quartzite pebbles, and the migration of large bed forms (sand waves/dunes). The finingupward transition from massive or cross-laminated grainstones to rippled fine-grained grainstones, and finally to muddy drapes colonized by oysters and platy corals suggests waning energy conditions

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Fig. 5. Logged stratigraphic section in the Ranero locality showing main lithologies, unconformities, distribution of age-diagnostic benthic foraminifers and cement stratigraphy. Age-diagnostic fossils after Lo´pez-Horgue et al. (2000). Black dots represent the presence of a particular cement zone for each analysed sample. Samples are located in their stratigraphic position within the lithological column.

during deposition and points to short-lived depositional events, with deposition of argillaceous material by suspension and benthic colonization during quiet periods. Physical correlation of sections in the down-dip direction shows terrigenous-rich lithofacies basinward of the coral-rudist buildups and sand shoal facies associations (Fig. 9). They are represented by fine-grained, terrigenous-rich calcarenites and sandstones with orbitolinids, and nodular to irregular bedded, bioturbated marls and marly limestones with marly, coral-rich beds. These facies suggest lower energy conditions than the sand shoal facies complex or slightly deeper conditions, most probably in an outer shelf environment.

Depositional sequences In this study, the early –late Albian succession of the Ramales area has been organized into four major depositional sequences (S1 to S4, Fig. 3d). Sequence boundaries are defined on the basis of

major unconformities and depositional facies shifts. At this time, the designation of these sequences is local. Sequence S1 is early Albian in age and comprises the upper part of the Ranero limestones. The top of sequence S1 in the Ojebar section (Figs 4 & 9) is a sharp, erosional surface covered by open-marine deposits of sequence S4. According to available biostratigraphy based on orbitolinids (R. W. Scott, pers. comm.), this surface represents a large hiatus spanning c. 5 Ma, from middle Albian to early late Albian (Lo´pez-Horgue et al. 2000; Garcı´a-Monde´jar et al. 2005). Detailed work carried out by Lo´pez-Horgue et al. (2000) documented karst features at this surface, suggesting sub-aerial exposure (Type 1 sequence boundary). Towards the platform margin, carbonates from sequences S2 and S3 onlap on this surface (Fig. 3d) and exhibit an angular unconformity with the underlying Ranero limestones (Lo´pez-Horgue et al. 2000). In the basin, sequences S2 and S3 form two wedge-shaped siliciclastic lithostromes that gradually filled the inherited depositional slope

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Fig. 6. Facies and microfacies of the lower Albian Ranero platform margin (see Fig. 5 for location of the samples). (a) Sample RAN-20. Field photograph showing a cross-bedded orbitolinid grainstone bar of the Ranero platform margin (Ranero section). (b) Sample RAN-20. Photomicrograph of the previous orbitolinid grainstone cemented by pore-lining marine fibrous cement and pore-filling equant spar (Ranero section). (c) Sample RAN-D. Photomicrograph of coral-skeletal grainstones of patch-reef. (d) Sample RAN-28. Bioclastic grainstones from the top of the Ranero unit showing early meniscus micritic cements (arrows).

(Fig. 3). Final infilling of the basin occurred at the top of sequence S3, which led to the deposition of shallow-water reefal carbonates on previous slope areas (Sopen˜a I limestones subunit, Figs 3 & 9). During deposition of these two infilling sequences, the Ranero limestones underwent erosion and subaerial exposure (Lo´pez-Horgue 2000). Sequence S4 corresponds to the upper part of the Sopen˜a limestones (Sopen˜a II subunit, Figs 3d & 9). During the deposition of S4, shallow-water conditions were re-established on the top of the Ranero limestones, and carbonates of this sequence rest unconformably on previous platform areas (Lo´pezHorgue et al. 2000). In the basin, the removal of the depositional relief between platform and basin allowed the growth of shallow water platform carbonates on former basinal areas (Fig. 3d). Vertical and lateral facies correlations reveal a retrogradational stacking pattern of depositional facies within these carbonates (Fig. 9) that indicates the effects of rapid flooding and platform back-stepping prior to the burial of the platform by siliciclastics of the Valmaseda Fm.

Petrography and diagenetic processes The studied carbonate rocks have been affected by a series of diagenetic processes that include at least six carbonate cement sequences, four phases of dissolution and several fracture generations. Cementation and dissolution are the main diagenetic processes affecting all the analysed samples. Carbonate cements appear as both replacement of skeletal grains (neomorphic grains), and as pore-lining and pore-filling cements in primary cavities, intergranular pores, dissolutional secondary porosity and fractures. The four representative localities logged in detail for description of depositional facies (Fig. 9), have also been investigated to study variations in the distribution of diagenetic products as interpreted from petrographic relationships. The location of the samples within each section is presented in Figures 4, 5, 7 and 10. Not all cement types are present in each sample, since their distribution is largely controlled by depositional facies, position of the sample with respect to the main unconformities (which controlled whether or

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Fig. 7. Logged stratigraphic section in Fuente Frı´a illustrating lithologies, main unconformities, and distribution of age-diagnostic benthic foraminifers and cement zones. Age-diagnostic fossils after Lo´pez-Horgue et al. (1994). Black dots represent the location of the studied samples and the presence of the particular cement zones for each analysed sample.

not early mouldic porosity and meteoric cements developed), and the pore size, with bigger pores showing a more complete sequence of cements.

Carbonate cement sequences and CL cement zones Six main sequences of cements have been recognized from CL analysis. A sequence of cements may be defined as a series of successive luminescent cement zones growing in continuity and limited from other sequences of cements by distinctive cement boundaries (Goldstein 1988; Braithwaite 1993). We have recognized two types of cement boundaries in the studied samples: (a) irregular trace boundaries that cut older cement zones and that are interpreted as recording dissolution prior to precipitation of the subsequent cement zone; and (b) zone boundaries characterized by crystallographic discontinuities, re-nucleation, or a change in the crystal growth orientation. Within each cement sequence, zones are always organized in the same

order, although one or more zones may be missing resulting in incomplete sequences. Not all cement sequences and zones are present in every sample and their relationships to fracturing and dissolution episodes are complex. Sequence 1 of cements (Z0). This is the oldest cement observed in some samples. It is characterized by isopachous, dusty, fibrous-to-bladed calcite crusts around skeletal grains that predate mechanical compaction and most of the mouldic porosity (Fig. 6b). They stain pink (non-ferroan) and fill primary intergranular porosity of grainstones from the platform margin of the Ranero limestones, as well as some secondary mouldic porosity and interconnected cavities within the reefal facies of the Sopen˜a limestones. Where present, these cements constitute 3.5–31.3% of the rock volume. Under CL, these cements show a patchy luminescence (cement zone Z0), indicative of recrystallization after precipitation. Fluid and solid inclusions are the primary cause of the dirty appearance

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Fig. 8. Facies and microfacies of the Upper Albian Sopen˜a limestones (see Figs 4 & 5 for location of the samples). (a) Ranero section, sample RAN-56. Organic mound overlain by deeper-water marls (incipient drowning episode, Sopen˜a II subunit). (b) Fine-grained, rippled calcarenites with marly flasers (Sopen˜a II subunit). (c) Sample RAN-60. Photomicrograph of fine-grained peloidal-skeletal grainstones of the previous rippled calcarenites. (d) Field photograph showing a grainstone shoal sequence from the Sopen˜a II subunit (Ojebar section, sample RA-305 D). The sequence shows an upward transition from massive, coarse-grained grainstone to wavy, medium-to-fine grained grainstones with mud drapes. Scale bar is 1 m. (e) Oolitic-skeletal grainstones from the previous sand bar (Sopen˜a II subunit, Ojebar section, sample RA-305 D). (f) Isolated quartzite pebble, 10 cm long, at the base of the grainstone bar of picture (e).

of these cements. Most solid inclusions are very small and unidentifiable, but some identified inclusions are relicts of iron oxides. Sequence 2 of cements (Z1 –Z3). Commonly occurs as scattered, non-ferroan, low-Mg scalenohedral calcite crystals (dogtooth spar) around pores, as syntaxial calcite overgrowths growing in optical continuity with echinoderm plates, or as continuous

pore-lining mosaic calcite cements rich in Fe-oxide solid inclusions. This cement developed in intergranular, shelter, intraskeletal and mouldic pores after leaching of aragonitic skeletons beneath the unconformities at the top of the Ranero and Sopen˜a I limestones. However, it has not been observed in samples from the Sopen˜a II limestones. This cement sequence is subdivided into three cement zones (Z1 to Z3, Fig. 11a, b).

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Fig. 9. Lithofacies stratigraphic distribution and facies panel reconstructed from correlation of measured sections and mapping of individual beds and surfaces.

The Z1 calcite zone is non-luminescent. Z2 shows an orange to bright yellow luminescence and may occur as a single thin yellow luminescent zone, or as an alternation of several thin yellow and nonluminescent/dull sub-zones. The Z3 zone is dully luminescent and may be sub-zoned. This sequence of cements postdates the earliest fibrous-to-bladed circumgranular cements (Z0) and most of the secondary mouldic porosity after aragonite dissolution, but it is interpreted as a relatively early cement because the embedded grains are usually uncompacted. Some samples just beneath the unconformity at the top of the Ranero limestones show meniscus and pendant micritic cements predating the Z1 –Z3 cement sequence (Fig. 6d). Sequence 3 of cements (Z4). This cement sequence is common in many samples and corresponds with clear, low-ferroan, inclusion-free, anhedral to euhedral equant spar calcite and some syntaxial overgrowths. It is a pore-lining or pore-filling mosaic calcite cement (depending on pore size), with crystal size increasing toward pore centres. This

spar cement has been observed partially to completely filling intergranular pores, fractures, and mouldic and vuggy porosity, and represents between 10 –32.5% of the rock volume. Precipitation of this cement occurred after a period of dissolution that affected some corals and skeletal grains. Under CL this cement is dark brown dull luminescent and corresponds with the calcite zone Z4. This is the first cement phase in samples from the Sopen˜a II limestones (Fig. 11c, d), and the second cement phase in samples from the lower Sopen˜a I limestones (Fig. 11a, b) and some samples from the platform margin of the Ranero limestones (Fig. 11e, f ). Z4 calcite may show a non-concentric compositional zoning of dull luminescence with non-luminescent to very dark dull irregular sectors (sector zoning, Fig. 11g, h). This particular pattern of CL is usually related to trace element partitioning and represents compositional differences between sectors that grew simultaneously (Reeder & Paquette 1989; Reeder 1991). Z4 can also lack sector zoning and be homogeneously dull, or in some well-developed syntaxial

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Fig. 10. Distribution of the oldest CL cement zones (from Figs 4, 5 & 7) superimposed on the lithostratigraphic cross-section of Figure 9. Black dots are the location of analysed samples. The earliest meteoric Z1–Z3 cement zones developed only beneath the two unconformities at top of the Ranero and Sopen˜a I limestones. Micritic limestones of the Ranero platform interior lack early meteoric Z1– Z3 cements. Z1–Z3 cements are interpreted to be younger in the Sopen˜a I limestones than in the Ranero limestones. Z4 is the first cement zone in the Sopen˜a II limestones.

overgrowths may display concentric dull and very dark dull sub-zoning. Minor fracturing (the first generation of fractures observed) and collapse of grains and micritic envelopes postdates this phase of cementation, suggesting a relatively early, shallow burial origin. Sequence 4 of cements (Z5–Z7). This cement occurs as blocky to poikilotopic, large (up to 5 mm in size) rhombic anhedral to euhedral crystals or as a replacement of former bioclasts or micrite matrix. It is Fe-rich and may contain few to abundant inclusions. The inclusions are generally out of optical microscope resolution but appear to contain fluid inclusions, remains of hydrocarbons (pyrobitumen) and pyrite. This cement constitutes from 0–37% of the rock volume and may occur as both pore-lining and porefilling cement, depending on pore size, but commonly completely occludes the remaining pore space in primary cavities, secondary pores and fractures (Fig. 12a –c). It precipitated after a new period of fracturing and dissolution, postdates mechanical and grain-to-grain compaction, but predates burialrelated stylolites Three cement zones (Z5 to Z7) are recognized within this cement sequence (Fig. 12a –d). Z5 is clear brown to dull orange

luminescent. Gradational sub-zonation due to minor colour variations occurs in several samples. Z6 luminesces dull brown, and Z7 is sub-zoned dull brown to dull orange. The later calcite zone may hold rare bituminous inclusions (Fig. 12e, f ) that indicate near-synchronous oil migration. This cement sequence was the first cement that extensively affected the micritic limestones of the Ranero platform interior along fractures formed before or contemporaneously to this cement, which are often solution-enlarged and filled by Z5 to Z7 cement zones (Fig. 12a, b). Sequence 5 of cements (Z8 –Z11). This sequence of cements occurs as the first carbonate cement in some secondary, non-fabric-selective solution cavities, large hydraulic breccias and large-scale extensional fractures oriented NW–SE. This suggests that vertical movement of fluids took place through these conduits. At least two generations of dolomite (Z8 and Z10) and two generations of calcite cements (Z9 and Z11) are recognized in thin section (Fig. 12f–h) by cross-cutting relationships and CL zonation, but there may be more in larger cavities. This sequence of cements postdates Z5 –Z7 ferroan calcite cements but still predates burial stylolites and later fractures.

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Fig. 11. Photomicrographs of paired plane-light (left) and CL (right) images (see Figs 4, 5 & 7 for location of the samples). (a)–(b) Sample FF-2, Sopen˜a I limestones, Fuente Frı´a section. Intergranular porosity is filled by Z1–Z3 (1– 3) pore-lining cements and Z4 (4) pore-filling cement. (c) – (d) Sample FF-5, Sopen˜a II subunit, Fuente Frı´a section. Note the lack of Z1–Z3 cement zones. Porosity is filled first by dark dull pore-lining Z4 (4) cement and later by orange dull Z5 (5) pore-filling cement. (e) – (f) Sample RAN-15, Ranero section. Neomorphic coral showing dissolution porosity filled by Z1–Z3 (1–3) pore-lining cement zones and pore-filling Z4 (4) cement zone. (g)–(h) Sample RA-305 D, Sopen˜a II limestones, Ojebar section. Intergranular and mouldic porosity is filled by Z4 cement (4), which shows sector zoning.

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Fig. 12. (a) – (b) Photomicrographs of paired plane-light (a) and CL (b) images. Solution enlarged fracture filled by Z5– Z7 (5– 7) cement zones from Ranero limestones in the Ojebar section (Sample RAN-45). (c) Intergranular primary porosity in grainstones of the Ranero platform margin filled by Z1–Z5 calcite zones (1– 5) (Sample RAN-10, Ranero section). (d) CL photomicrograph of Z11 orange-bright calcite (11) filling irregular fissures and dissolutions on previous Z5–Z6 burial calcites (5– 6) (Sample FF-30, Fuente Frı´a section). (e) Pyrobitumen inclusions (B) in burial Z7 sparry calcite cements (Sample RAN-D, Ranero Section). (f) Pyrobitumen-rich (B) Z7 calcite (7) and hydrothermal Z8 saddle dolomite (D, 8) (Sample RAN-D, Ranero Section). (g)–(h) Paired plane-light (g) and CL photomicrographs (h) of hydrothermal Z9 calcite (9) and saddle Z10 dolomite (10). In CL saddle dolomite is red and shows dissolutional contacts (arrows) with the dull unzoned Z9 calcite (Sample RAN-E, Ranero section).

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Macroscopically, both replacive and pore-filling dolomites and calcites exhibit a patchy distribution as metre- to kilometre-scale, irregular masses and fracture-fills. Zone Z8 consists of coarse saddle (baroque) dolomite cement with dark inclusions (Fig. 12f ). These dolomites show coarse anhedral-spar to euhedral crystals, the latter with undulose extinction and curved faces and cleavage, and are dull luminescent. Zone Z9 corresponds with coarse rhombohedral crystals of spar calcite greater than 1 cm in size. It is low-ferroan to ferroan and shows a non-concentric dull brown and dull clear zonation (Fig. 12g, h). This calcite postdates all previously reported calcite cements but predates some saddle dolomite (Z10) and still predates stylolites. Z10 is red under CL and corresponds to a younger stage of ferroan saddle dolomite (Fig. 12g, h). The contact between Z9 calcite and Z10 dolomite locally is a dissolution surface (Fig. 12g, h). This dolomite fills open porosity in hydraulic breccias, cavities and fractures that cross cut both carbonate matrix and all previous cements. Locally, dolomites may be replaced by calcite (dedolomitization), which is corroded and cloudy due to Fe-oxide inclusions. Z11 is one of the last calcite cements and precipitated after minor dissolution that affected previous cements. It has a dull orange to bright orange luminescence (Fig. 12d). Sequence 6 of cements (Z12). This sequence of cements consists of sparry, low-ferroan calcite cement and has been identified as the Z12 calcite zone. This is a later stage of dull luminescent calcites filling tectonic fractures and veins that crosscut saddle dolomite, coarse calcites and stylolites.

Other diagenetic products Pyrite occurs as small cubes or massive aggregates that postdate most of the calcite cement phases, suggesting a late stage of precipitation. Authigenic quartz occurs as euhedral crystals with abundant calcite inclusions. Silica replaces micritic matrix, carbonate grains and sparry calcite cements, suggesting silica precipitation following carbonate cementation of the host. Furthermore, although absent in the studied samples, the study area hosts abundant Pb –Zn ore deposits (sphalerite, galena, fluorite, barite) in the dolomitic patches and in later veins. Based on paragenetic relationships and microthermometric and REE studies of fluorites, these mineral deposits have been associated with a later hydrothermal fluid circulation probably related to the Alpine orogeny (Herrero et al. 1988). Goethite and Fe-oxides are opaque and well distinguished by the typical red colour in reflected light. Petrographic relations suggest that they formed

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by the replacement of pyrite and ferroan cements within an oxidizing diagenetic environment, probably during the Alpine uplift.

Stages of porosity development Present-day open porosity from modal analysis varies from 0 –1.5%, and includes rare moulds of leached corals and isolated vugs that have not been completely occluded by cements, intercrystal voids in patchy dolomites, and rare open fractures and stylolites. Primary intergranular porosity in grainstone, as well as primary skeletal and shelter cavities in reefal lithofacies has been completely occluded by fibrous-to-bladed and blocky, Z0 to Z7 calcite cements. Secondary porosity development was important and included dissolution porosity (0.7–53%) as mouldic, solution enlarged mouldic, channel and vuggy porosity, and fractures (up to 7.7%). Dissolutional secondary porosity developed at different stages during shallow to deep burial realms. At least four stages have been recognized. Dissolution stage I. This dissolution stage corresponds with early pre-compaction leaching of aragonitic skeletal grains (mainly corals, gastropods and molluscs). The occurrence of this fabric-selective mouldic porosity indicates that the dissolving fluid was undersaturated with respect to aragonite, but supersaturated or in equilibrium with calcite. Moulds formed during this dissolution stage were partially filled with Z1 to Z3 calcite cements (Fig. 10e, f ). This dissolution event affected carbonates of the Ranero platform margin and the Sopen˜a I limestones. The lack of these cement zones in moulds of the Sopen˜a II limestones suggests that dissolution I occurred before its deposition. Dissolution Stage II. This second stage of secondary porosity developed shortly after deposition of the Sopen˜a II limestones. It predates precipitation of Z4 calcite and occurred before or during mechanical compaction of the Sopen˜a II carbonates. Dissolution during this stage mainly affected metastable components of the Sopen˜a II limestones, resulting in moulds filled with Z4 calcite cement (Fig. 11g, h). Dissolution Stage III. More aggressive than the previous dissolution stages, dissolution stage III affected aragonitic skeletal grains but also micritic matrix creating vugs and enlarged mouldic and fracture porosity. This generation of secondary porosity developed before precipitation of the ferroan sparry calcite of cement sequence 4 (Z5–Z7), and postdates neomorphism, mechanical compaction, and some fracturing and spar calcite (Z1– Z4). Microscopic collapse of grains and microspar geopetal infilling in vugs is common. Ferroan Z5 to Z7

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Fig. 13. Geohistory and organic matter maturity of the Ramales area using Platte River’s BasinMod (1D modelling). (a) Relationship between syn-rift and post-rift subsidence and heat flow in an extensional basin (based on McKenzie 1978). Two peaks of maximum heat flow during the late Jurassic– earliest Cretaceous and Early Cretaceous (Albian) rifting pulses have been modelled as an initial working hypothesis. (b) Vitrinite reflectance values (%Ro) of the main Mesozoic stratigraphic units of the Ramales area plotted v. depth. Horizontal lines around the values represent error bars. The profile is not sub-linear suggesting higher thermal maturity than that predicted from rates of subsidence and sedimentary thicknesses alone (sub-linear blue line), and the assumed heat flow from Figure 13a. (c) New calculation of the relationship between syn-rift and post-rift subsidence and heat flow for the BasqueCantabrian basin. Assumptions incorporated in this new calculation include a thermal anomaly during the late Albian– Late Cretaceous. (d) Modelled vitrinite reflectance values (blue line) from the subsidence and heat flow relationship assumed in Figure 13c. The new profile fits well with the measured Ro values. (e) Modelled Ro profile v. depth and geologic time calibrated with the real measured Ro values. Stratigraphic units were sequentially back-stripped and decompacted. The isoreflectance values used to define organic matter maturity through time are represented by the colour scale. (f) Modelled temperature v. depth and geologic time for the Ramales area. Palaeoisotherms were

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carbonate cement partially to completely occludes this generation of secondary porosity. This phase of secondary porosity extensively affected the micritic limestones of the Ranero platform interior, which were almost unaffected by the previous dissolution stages I and II. Dissolution Stage IV. This dissolution event is related to the saddle dolomite and is mostly restricted to platform margin areas and near fault zones. Dissolution during this stage resulted in both large-scale solution collapse and hydraulic breccia bodies and micro-scale, non fabric-selective solution cavities filled by coarse spar calcites and saddle dolomites (Z8 to Z11 cement zones, Fig. 11d, g, h). Fracture porosity. We have recognized at least five generations of fractures: three pre-stylolite generations, one stylolite-related generation, and at least one generation postdating stylolites. Pre-stylolite fractures are assumed to have been formed during periods of extensional tectonic activity and/or by sedimentary loading. The first recognized generation of fractures (I) cross-cuts Z1 –Z3 calcite zones and was filled by Z4 calcite cements. The second generation of fractures formed after the Z4 calcite zone but before dissolution stage II and before Z5 calcite cement. Fractures formed during this stage are generally solution-enlarged and are filled by Z5– Z7 cement zones. Fractures of the third generation (III) cross-cut late calcite cements, but still predate stylolites. They are filled by saddle dolomite and coarse sparry calcite (Z9 and Z10). Stylolite-related fractures are associated with compressional stress. They display an irregular, discontinuous to anastomosed pattern, from a few microns up to 5 mm wide, and are filled by late blocky calcite (Z12). Post-stylolite fractures cross-cut all the previous diagenetic features, including stylolites and Z11 bright calcites and are infilled by dull calcite cements (Z12 or later).

Subsidence and thermal modelling In this study, the Mesozoic –Neogene subsidence history and one dimensional thermal modelling of the Ramales area have been estimated by using the back-stripping methodology and BasinMod software (Fig. 13). Across unconformities, stratigraphic

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thicknesses were estimated and restored, and then removed to reflect the influence of tectonics on the burial history. Maximum loading of sediments before Alpine compression is difficult to calculate due to the erosive removal of overlying Upper Cretaceous and Cenozoic sediments. An approximation of their thicknesses has been estimated by comparison with the closest outcrops (within few kilometres of the modelling location) which indicate loading by Albian carbonates of at least 3000 m of sediment. A present day geothermal gradient of 25 8C km21 was inferred from formation temperatures obtained from nearby industry boreholes. The heat flow history of a basin is mostly controlled by the lithospheric mechanics and amount of stretching involved in basin development, assuming the thermal conductivity is known (Allen & Allen 1990). In an extensional basin, the total subsidence and associated heat flow can be quantified by the McKenzie (1978) model, which considers an initial fault-controlled subsidence, modelled as instantaneous, and a subsequent thermal subsidence that decreases exponentially with time. To quantify the heat flow in this study, the McKenzie (1978) model has been applied to the post-Triassic succession using the BasinMod heat flow options and assuming, as an initial hypothesis, two extensional pulses during the Late Jurassic– Early Cretaceous and early Albian respectively (Martı´n-Chivelet et al. 2002) with palaeo-heatflow peaks of c. 75 mW m22 (Fig. 13a). A theoretical maturity profile as a function of depth has been produced by this modelling (Fig. 13b). To determine whether these geothermal values are reasonable Mesozoic estimates, the model has been calibrated using thermal maturity values (vitrinite reflectance, Ro) measured from Lower Jurassic to Upper Cretaceous samples that outcrop in the vicinity of the study area. A plot of Ro values against formation depth indicates that maturity values for pre-Late Cretaceous deposits depart from the modelled values (Fig. 13b). This indicates that these rock units were exposed to thermal conditions higher than those predicted by the model and suggests a thermal perturbation during the burial history. Moreover, illite crystallinity indices of Lower –Middle Cretaceous shales from the central BCB also indicate anomalously high temperatures (Aro´stegui et al. 1991). In the axial zones of the BCB, volcanism began during

Fig. 13. (Continued ) calculated using the kinetic maturation model of Sweeney & Burnham (1989). The diagram shows a peak heat flow in the latest Albian– Cenomanian (thermal anomaly), after which heating. Stratigraphic units are 1, Triassic continental deposits and evaporites; 2, Lower Jurassic marine deposits and hydrocarbon source rocks; 3, Upper Jurassic–Lower Cretaceous continental and transitional deposits; 4, Lower Cretaceous continental and shallow marine siliciclastics; 5, Aptian platform carbonates; 6, lower Albian platform carbonates; 7, upper Albian platform carbonates; 8, upper Albian– lower Cenomanian platform siliciclastics (Valmaseda Fm); 9, Upper Cretaceous marine carbonates.

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the late Albian –Cenomanian (Mathey 1986) in response to maximum lithosphere stretching, high tectonic subsidence and consequent high sedimentation rates (Garcı´a-Monde´jar et al. 1996). This may have induced heat flows of c. 100 mW m22 on average during this time (Allen & Allen 1990). Applying heat flow conditions associated with two extensional pulses and a thermal anomaly of c. 100 mW m22 at the beginning of the Late Cretaceous (Fig. 13c), the model yield calculated Ro values that are consistent with the measured values (Fig. 13d). This indicates that our model is properly calibrated. The modelled maturity and thermal histories of the Ramales area are shown in Figure 13e and f. Temperature palaeoisotherms suggest that the Albian carbonate rocks were exposed to maximum temperatures of c. 150 8C, and that these temperatures were reached at c. 92 Ma (Cenomanian– Turonian boundary) and at a burial depth of about 1700–2200 m (Fig. 13f ). Primary fluid inclusion data from hydrothermal calcites from the study area (Simon et al. 1999) indicate homogenization temperatures of 120 –160 8C, and mesosalinities of 14–20 wt% NaCl equivalent, which are in accordance with those predicted from the thermal model. Lead isotope values of galena samples from different mineral deposits hosted in Urgonian limestones of the BCB indicate ages for hydrothermal mineralization in the basin of c. 115 –60 Ma (Velasco et al. 1996, 2003), which are consistent with those predicted by the thermal modelling.

Stable isotopes Results The sequence of cements 1 to 6 were analysed for their carbon and oxygen isotopic composition and their values are shown in Figure 14. Samples from well-preserved rudist shells and brachiopods are included for comparison to Albian C and O isotope composition of seawater precipitates in the BCB. The d13C values are between 21.3‰ and þ3.7‰ PDB, and the d18O values are between 22.4‰ and 215.7‰ PDB (Fig. 14a). Cement sequence 1 (Z0) and skeletal marine components have similar d13C and d18O values, varying, respectively, between þ0.33‰ and þ3.7‰ PDB and between 22.4‰ and 22.9‰ PDB. Spar Z1 to Z3 calcite (cement sequence 2) has d13C values ranging from 21.3‰ to þ2.7‰ PDB, and d18O values varying between 24.3‰ and 27.5‰ PDB. The isotopic composition of cement Z4 (cement sequence 3) range between þ0.8‰ and þ1.5‰ PDB for carbon and between 27.9‰ and 28.7‰ PDB for oxygen. The d13C values of cement sequence 4 (Z5 –Z7) are between þ0.8‰ and

Fig. 14. (a) Stable isotope composition of different components and cements of the Ramales carbonates. (b) Equilibrium relationship between d18O of water, temperature and d18O of calcite from the Kim & O’Neil (1997) equation [21000 lna(calcite-water) ¼ 18.03 (103 T21) 2 32.42]. It shows the estimated range of d18OSMOW and temperature of seawater and the estimated local d18OSMOW of meteoric water and basinal brines.

þ3.1‰ PDB, whereas Z5 to Z7 cements have d18O values ranging from 27‰ to 210.1‰ PDB. Within cement sequence 5, the Z9 calcite cement has d18O values ranging from 210.8‰ to 212.4‰ PDB and d13C values ranging from þ1.3‰ to þ1.8‰ PDB, whereas the saddle dolomite (Z10) has slightly less negative d18O values (29.1 and 29.5‰ PDB) and slightly more positive d13C values (þ2.3‰ and þ3.7‰ PDB). One sample of the latest sequence of cement 6 filling a late generation of fractures (Z12 calcite) has yielded a d18O

POROSITY, DIAGENESIS AND BASIN MODELLING

value of –15.7‰ PDB and a d13C value of þ0.8‰ PDB (Fig. 14a).

Interpretation The carbon and oxygen stable isotope compositions of the main cement sequences show relatively constant d13C values but a wide range of d18O values. Compositional fields show a progressive displacement to more negative d18O values with decreased relative age of the cements and increased interpreted burial depth. This suggests a trend toward more elevated temperatures with burial, with carbon derived from the host rock. Except for one sample (one Z1–Z3 cement sample with d13C of –1.3‰ PDB), the d13C values of all analysed calcite and dolomite cements (0.7‰ to 3.7‰ PDB) overlap those of the biogenic and abiotic marine calcite (0.3‰ to 3.7‰ PDB). This suggests that the d13C composition of pore waters was buffered by the host carbonates, with minimal influence from CO2 derived from other sources such as soil-gas or maturation of organic matter. The oxygen isotopic composition of carbonate precipitates is temperature and salinity dependent, with d18O values of calcite usually decreasing with increasing temperature and decreasing salinity. Data from the Ramales platform show a progressive decrease in their d18O values from older to younger cement sequences (Fig. 14a), suggesting a decrease in salinity of diagenetic fluid and/or increase in temperature with increasing burial depth. Mean d18O values of skeletal components and Z0 isopachous calcite cements are very similar (22.8 and 22.4‰ PDB respectively). This isopachous cement is similar in texture and isotopic composition to fibrous marine cement described in carbonates from many different places and ages (James & Choquette 1983). This suggests a marine (seafloor) origin for the earliest fibrous-to-bladed calcite crusts. Z1–Z3 spar calcite cements have more negative d18O values (24.3 to 27.5‰ PDB). In other studies, similar blocky cements have been interpreted as being of either marine-burial or meteoric origin (Melim et al. 2002). However, spar calcite precipitated from marine-burial diagenesis usually has more positive oxygen stable isotopic values, which indicates precipitation from normal to slightly modified marine pore fluids (Melim et al. 2002; Ehrenberg et al. 2006). The more negative d18O values of Z1 –Z3 spar calcite relative to marine cements (Z0) and skeletal calcite are compatible with precipitation in a shallow meteoric aquifer (Fig. 14b). The d18O composition of calcite precipitated from modern meteoric groundwater in tropical areas with average monthly temperatures of c. 20 8C is about –4‰ (+2) PDB, whereas d13C values may

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show a large range (Cerling 1984; Lohmann 1988). The positive to slightly negative d13C values of the Z1 –Z3 cements, if meteoric in origin, suggest low relative amounts of CO2 derived from plant respiration or oxidation of vegetal material in the soil zone. This may reflect either poor soil development on the sub-aerially exposed platform top, or a very high contribution of CO2 derived from carbonate dissolution (Muchez et al. 1998; Wu & Chafetz 2002). The latter is the most likely, given that this cement developed between extensive dissolution related to stages I and II. The d18O values of Z4 calcite (27.9 to 28.7‰ PDB) overlap or are more negative than the d18O values of the Z1 –Z3 cements (Fig. 14a), but still plot within the field of meteoric precipitates reported elsewhere (e.g. Lohmann 1988; Moore 1989; Muchez et al. 1998). Z5– Z7 cements have similar to more negative d18O values (27‰ to 210.1‰ PDB) than previous cement sequences. Z9 cements have more negative d18O values (210.8‰ to 212.4‰ PDB) than Z5–Z7 cements, whereas saddle dolomites (Z10) have slightly less negative d18O values (29.1 and 29.5‰ PDB). This trend probably reflects a combination of varying fluid isotopic composition and/or fluid temperatures as well as a difference in fractionation between calcite and dolomite that precipitate from the same fluid. Dolomite and calcite precipitated from the same fluid usually exhibit different oxygen isotopic values, with dolomite being more positive by c. 3‰ (Land 1980). Finally, the d18O value (215.7‰ PDB) of one sample of calcite filling a late generation of fractures (Z12 calcites) is in agreement with precipitation from deep burial, hot basinal brines (Moore 1989).

Timing and source of the diagenetic fluids The different dissolution and cement stages that affected the Albian carbonates at Ramales formed in a variety of diagenetic environments (Fig. 15) that can be interpreted in the context of the regional burial and uplift history of the Ramales platform (Fig. 16). The thermal model of the study area (Fig. 13f ) records a heating and cooling history that follows the pattern of this burial history. The thermal model displays a heat flow event at c. 92 Ma, roughly coincident with the Cenomanian –Turonian boundary (Fig. 16) when the Albian carbonates were buried to about 1700– 2200 m and heated to 150 8C (Fig. 13f ). Therefore, most of dissolution and precipitation of calcite and dolomite cements took place at or below 150 8C. If we accept that cement sequences precipitated in equilibrium with the ambient fluids, and using

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Fig. 15. Summary of the paragenetic sequence inferred from this study for the Albian platform carbonates of the Ramales area. Chart shows the relative timing of carbonate stabilization, cements and diagenetic events. Diagenetic processes have been separated into those that created, destroyed or had little effect on porosity development.

the burial temperature range for the studied carbonate succession that has been deduced from the thermal model (Fig. 13). If we take the oxygen isotopic relationship for calcitewater given by the expression of Kim & O’Neil (1997): 1000 lna(calcitewater) ¼ 18:03(103 T1 )  32:42 we can use the d18O values for calcite precipitated at equilibrium to estimate the O-isotopic range and possible source of fluids that might have precipitated the different cement sequences (Fig. 14b). Fig. 16. Burial history of the studied Albian platform carbonates of Ramales deduced from 1D modelling, with suggested timing of the main paragenetic events. Horizontal lines represent approximate boundaries between diagenetic stages.

Seafloor marine diagenesis. Seafloor marine diagenesis is characterized by widespread micritization of grains, micritic envelopes, and circumgranular

POROSITY, DIAGENESIS AND BASIN MODELLING

fibrous high-Mg calcite cements. Most interparticle pores and shelter cavities in skeletal grainstones and reefal limestones of the Ranero platform margin were lined with fibrous-to-bladed calcite (sequence 1 of cement, Z0). This cement (Z0) has all the attributes of marine precipitates at the seafloor but their presence within some moulds of leached aragonitic grains is problematic and could suggest diagenesis by mixing-zone or meteoric fluids (e.g. Halley & Harris 1979; Melim et al. 2002). However, since Cretaceous seawater was probably undersaturated with respect to aragonite (Sandberg 1983; Stanley & Hardie 1998), some dissolution of aragonite may also have occurred in skeletal grains exposed for a long time on the sea floor (Palmer et al. 1988). Petrographic and CL analyses indicate that these precipitates are neomorphic. These cements suggest an origin as high-magnesium calcite (HMC), which was later transformed during neomorphism to lowmagnesium calcite (LMC) with little loss of textural details (Lohmann & Meyers 1977). Their stable isotope composition (d13C of þ2‰ and d18O of –2.4‰ PDB) is comparable to other reported stable isotope values of Cretaceous tropical marine cements (Pearson et al. 2001). This suggests that stabilization of the HMC to LMC occurred in a closed system in contact with marine pore waters, most probably shortly after burial. Therefore, we can use the d18O values for calcite precipitated at equilibrium to calculate palaeotemperatures using estimates for d18O of contemporaneous seawater and the expression of Kim & O’Neil (1997). Using a d18OSMOW value of –1‰ (Pearson et al. 2001) and a mean d18O value of – 2.4‰ PDB for the marine cements, a local temperature of about 25 8C for the tropical surface seawater is obtained (Fig. 14b).

Meteoric phreatic diagenesis Using the Kim & O’Neil (1997) expression (Fig. 14b) and our palaeotemperature estimate for ambient surface water (25 8C), sequence 2 of cements (Z1 –Z3), with d18O values between –4.3‰ and –7.5‰ PDB, probably precipitated from a fluid with a d18OSMOW value of about – 2.3‰ to –5‰. The range of these d18OSMOW values is interpreted to indicate precipitation from meteoric water (e.g. Suchy et al. 2000; Fu et al. 2008). Meteoric phreatic diagenesis was probably related to the development of meteoric lenses beneath two main regional unconformities. These are the middle–earliest late Albian hiatal unconformity (c. 5 Ma) that affected carbonates of the Ranero unit, and the late Albian (auritus sub-zone) unconformity that affected carbonates of the lower Sopen˜a I subunit (Fig. 3). Field evidence suggests

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that these unconformities are associated with Type 1 sequence boundaries controlled by a combination of relative sea-level changes and tectonic influence namely block faulting related to rifting pulses (Lo´pez-Horgue 2000; Garcı´a-Monde´jar et al. 2005). Differential uplift created gradients in topography, which probably induced a topographically-driven influx of freshwater. Below these unconformities, grainstone and reefal facies with abundant aragonitic components (corals, rudists, molluscs) exhibit extensive mouldic, pre-compaction porosity (dissolution stage I) followed by precipitation of pore-lining Z1 to Z3 cement zones. The non-ferroan composition, the CL character, the isotopic values, and the presence of Fe-oxide solid inclusions for the Z1– Z3 calcites suggest an oxidizing to slightly reducing, near surface meteoric origin (Muchez et al. 1998). In the oxidizing environment, Fe2þ cannot be reduced and substituted for calcium in the calcite lattice, therefore all Fe2þ available in the pore water may have precipitated as inclusions of iron oxides. This process is well documented in the oxidizing portions of modern aquifers (Champ et al. 1979) and the distribution of the non-luminescent Z1 cements (Fig. 10) may define the distribution of oxidizing portions of the palaeoaquifer system (e.g. Mussman et al. 1988; Dickson & Saller 1995; Mutti 1995). The bright to dull colours of the Z2 and Z3 calcites typically reflect higher concentrations of Mn2þ and Fe2þ that cause luminescence (e.g. Meyers 1978; Grover & Read 1983). It is most likely that they formed under sub-oxic conditions following stagnation of the palaeoaquifer with burial or when the recharge of the palaeoaquifer finished (e.g. Meyers 1978; Mussman et al. 1988; Goldstein 1988; Niemann & Read 1988; Braithwaite 1993; Dickson & Saller 1995). Incursions of oxidizing meteoric water beneath the late Albian, auritus sub-zone unconformity penetrated to a depth of at least 100 m (depth of the last nonluminescent Z1 cement). In the Ranero limestones, early secondary porosity filled by Z1 to Z3 cement zones is common in limestones of the platform margin and the cements extended at least 200 m below the top unconformity. Surprisingly, these cements are absent in carbonates of the Ranero platform interior (Fig. 10), suggesting that the development of early meteoric diagenesis was faciesrelated. Since the platform margin facies are mainly made of coral (aragonite) limestones and packstones-grainstones, the result is a greater potential for fluid flow and diagenetic modification during sub-aerial exposure than in the micritic limestones of the platform interior which had lower primary porosity. As a consequence, platform margin grainstones and coral-reef facies display greater rates of early carbonate dissolution and subsequent meteoric

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cements during the development of meteoric lenses. An alternative explanation is that the unconfined coastal aquifers developed freshwater lenses with very local extension or that only the platform margin was sub-aerially exposed and therefore influenced by meteoric diagenesis during this stage, although this would require significant relief on margin morphology. Shallow burial-meteoric diagenesis. Shallow burial diagenesis was characterized by dissolution of metastable grains (dissolution stage II), low-Mg Z4 spar cement, mechanical compaction and some fracture generations (fracture generations I and II; Fig. 15). The petrography of the Z4 cement suggests either a meteoric or marine-burial origin for the pore fluid (Melim et al. 2002). The dull brown CL character suggests higher Fe concentrations and more reducing pore waters than the earlier cements, which characterize burial diagenetic environments. However, whereas shallow burial calcites precipitated from marine pore-waters usually have marine oxygen isotopic signatures (Melim et al. 2002), those from meteoric waters have more negative values. The negative oxygen isotope values of Z4 calcites (27.9‰ to 28.7‰ PDB) that equal or are slightly more negative than meteoric Z1– Z3 calcites suggest a buried meteoric pore fluid (Fig. 14b). Therefore, we interpreted this cement to have been precipitated in a shallow meteoricburial environment. The distribution of Z4 calcite as the first porelining cement in the Sopen˜a II limestones, as well as a pore-lining and pore-filling cement in the Ranero platform margin and Sopen˜a I stage, indicates that significant porosity remained open during shallow burial. The lack of dominantly oxidizing (non-luminescent) meteoric cement (Z1– Z3) in the Sopen˜a II unit points to an external source for the fluids, probably meteoric-derived groundwaters in the downflow portion of aquifers sourced from sub-aerially exposed platform sectors outside the study area. Fluid would have flowed laterally into the carbonates during deposition upon the platform of the Valmaseda shales, which acted as an aquiclude. The aquifer was confined at this time, allowing a gravity-driven flow of meteoric water from recharge areas located in uplifted areas probably to the NE (Rosales 1999). Recharge of meteoric ground waters to the Ramales aquifers would have ceased when the uplifted recharge areas were finally drowned and buried. By Cenomanian times, all of the surrounding Albian carbonate outcrops were buried beneath the marine siliciclastics of the Valmaseda Fm and their lateral equivalents (Garcı´a-Monde´jar 1990), and meteoric recharge would have ceased, by the proposed mechanism.

Intermediate burial diagenesis This stage of diagenesis involved processes of dissolution (stage III) and cementation by sequence 4 (Z5 –Z7) ferroan calcites (Fig. 15). This cement sequence has been interpreted to form in intermediate burial depths given that they postdate all mechanical compaction but are truncated by burial-related stylolites. The coarse sparry crystals are probably the result of slow rates of growth related to restricted and slow fluid flow through the pores (Gonza´lez et al. 1992). In addition, the change from non-ferroan to ferroan cements is conventionally interpreted to reflect a change in the diagenetic environment from shallow burial to deeper burial (e.g. Choquette & James 1990). The elevated Fe content probably reflects allochthonous compactional water flowing laterally from siliciclastics that contain feldspars or other Fe-bearing minerals, which are easily dissolved during the diagenesis of sandstones laterally equivalent to the limestones. Such fluids have been considered as a main source of Fe for burial calcite cements in carbonates elsewhere (Moore 1989). Fluctuations in the available Fe2þ may be responsible for the dull orange-banded luminescent zones of these cements (Dorobek 1987; Reeder 1991). Dissolution features during this diagenetic stage include moderate to abundant amounts of enlarged mouldic porosity, vugs, solution-enlarged fractures and channel porosity that also affected the carbonates of the Ranero platform interior. Fractures formed before or contemporaneously to this dissolution stage were often solution-enlarged and are filled by Z5– Z7 cement zones. These cement zones are the first carbonate cements to have extensively affected the micritic carbonates of the Ranero platform interior (Fig. 10). It is possible that loading or tectonic fissures related to maximum basin stretching that occurred at this time (Garcı´aMonde´jar et al. 1996) developed vertical conduits in the Ranero platform interior and allowed fluids to penetrate into the otherwise tight lagoonal micritic limestones. This probably resulted in the development of solution-enlarged fractures that connected moulds (Fig. 12a, b). This is supported by the distribution of this cement that is restricted along solution enlarged fractures and moulds connected by fissure systems. Solution-enlarged moulds resulted in vuggy porosity, which presents abundant microscopic solution-collapse features that are extensively cemented by Z5– Z7 ferroan calcite. The d18O values for this calcite cement (27.1‰ to 210.1‰ PDB) precipitated at equilibrium from water and can be used to estimate the temperature of precipitation using the expression of Kim & O’Neil (1997). If we assume the fluid originated from formation waters, estimated fluid

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temperatures using the equation above with an initial d18O ratio of 21‰ for the Cretaceous seawater (Pearson et al. 2001) would range from 45– 65 8C. Assuming an average initial water temperature of about 25 8C for tropical seawater in the BCB and a mean geothermal gradient of 30 8C km21, these cements would have precipitated at burial depths of about 700 –1500 m. These depth estimates are consistent with those derived in the subsidence and thermal modelling for the late Cenomanian (Figs. 13f & 16). The occurrence of pyrite and hydrocarbon inclusions trapped in these cements suggests that some generation and release of hydrocarbons and H2S took place during this stage (Figs 12e, f, 15 & 16).

Hydrothermal stage The increase of burial diagenesis was accompanied by the emplacement of sequence 5 saddle dolomite and ferroan calcite cements (Z8–Z11) that occur within dissolution cavities and fractures generated shortly before the emplacement of these cements (dissolution stage IV). Sequence 5 calcites and dolomites have positive d13C values similar to that of the host limestones. This suggests that the CaCO3 released to the pore fluid by this dissolution phase provided the HCO2 3 for the precipitation of this new sequence of calcite and dolomite cements. Saddle dolomites require a minimum of 60–80 8C to form and the full temperature range is reported to be 60 –150 8C (Radke & Mathis 1980; Sibley & Gregg 1987). Previous fluid inclusion studies in the Ranero area suggest mesosaline fluids and temperatures as high as 120 –160 8C (mean 140 8C) for these cements (Simon et al. 1999). The low d18O composition of the associated calcite (210.8‰ to 212.4‰ PDB) also suggests high temperatures. If we combine available microthermometry from fluid inclusions (Simon et al. 1999) with the oxygen stable isotope data of the calcite, the calcite would have precipitated at equilibrium from a fluid with an original d18O composition of about 6–8‰ SMOW (Fig. 14b). These positive d18O values rule out a possible meteoric origin for the fluid (because of the strong negative values that typify meteoric fluids), and are much higher than d18O values of normal seawater. Therefore, the positive d18O values of the hydrothermal fluid likely reflect an origin from fluids enriched in 18O and having salinities higher than marine-derived fluids. This change to more elevated salinities may suggest that the fluids involved were basinal brines or alternatively derivation from a deep-seated source such as the underlying Triassic evaporites. Later burial and tectonic uplift. Deeper burial diagenesis is characterized by chemical compaction

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(stylolites), extensive fracturing and circulation of hot basinal fluids through tectonic fractures (Z12 calcite zone and later cements and mineral deposits). Finally, tectonic uplift during the Alpine orogeny may have resulted in dedolomitization and oxidation of Fe-carbonates and pyrite and the subsequent precipitation of Fe-oxides.

Conclusions Two major carbonate depositional systems developed at Ramales during the studied time interval: a lower Albian steep-sloping platform (Ranero unit) and an upper Albian low-gradient open shelf (Sopen˜a unit). Within these platform carbonates two unconformities have been recognized that can be traced from the platform interior to the basin: a hiatal unconformity spanning the middle – early late Albian (c. 5 Ma constrained with biostratigraphic data) that separates the Ranero and Sopen˜a limestones; and a later Albian (auritus sub-zone) unconformity, which separates the Sopen˜a limestones in two subunits (I and II). This study proposes that sub-aerial exposure occurred in association with these two unconformities. Petrographic, cathodoluminescence and geochemical analyses demonstrate that major pervasive meteoric dissolution and cementation occurred in strata below these main unconformities. This resulted in leaching of aragonitic components followed by precipitation of pre-compaction pore-lining Z1 –Z3 cement zones. The petrographic and stable isotopic composition of these cements suggest that the fluids were sourced from oxidizing to slightly reduced meteoric waters in unconfined aquifers during relative sea level lowstands. In the Ranero unit, meteoric diagenesis affected mainly grainstones and coral reefs in the platform margin facies but does not seem to have affected the platform interior micritic carbonates. The studied platform units have minimal present-day porosity but are interpreted to have been porous and permeable during deposition and shallow to intermediate burial by the creation of dissolution secondary porosity (dissolution stages I to III). Extensive cementation and replacement by meteoric (Z1 –Z4 cement zones) and basinal fluids (Z5 –Z7 cement zones) transformed the porous limestones into non-porous limestones by the occlusion of nearly all-remaining porosity prior to late burial. The platform was almost completely cemented during intermediate burial, with the bulk being reduced to less than 3% of rock volume. Additional secondary porosity was subsequently created (dissolution stage IV), which may have resulted from the expulsion of hot mesosaline basinal brines from underlying and adjacent formations during compaction or tectonic movements. This later

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porosity was partially occluded by the precipitation of several generations of very coarse ferroan calcite and saddle dolomite (Z8– Z11 zones). Carbonate diagenesis related to these fluids created localized dolomite geobodies and large-scale porosity associated with fault zones and fractures but did not significantly increase the small-scale porosity in the host limestones. This is a contribution to the IGME research project ‘275/ CANOA 35006: Relacio´n entre sedimentacio´n, tecto´nica y flujo de fluidos durante la extensio´n del Creta´cico inferior en la cuenca de Santander’. Funding for a preliminary research project was provided by YCI Espan˜a S. L. The authors are grateful to S. Quesada (Repsol Exploracio´n) who assisted with BasinMod software for subsidence and thermal modelling and kindly provided unpublished vitrinite reflectance data. Critical comments by reviewers O. Weidlich (Sultan Qaboos University, Oman), P. Burgess (Shell International) and editor K. D. Gerdes greatly improved the final version of the manuscript.

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Stratigraphic architecture and fracture-controlled dolomitization of the Cretaceous Khami and Bangestan groups: an outcrop case study, Zagros Mountains, Iran I. SHARP1*, P. GILLESPIE1, D. MORSALNEZHAD2, C. TABERNER3, R. KARPUZ4, J. VERGE´S5, A. HORBURY6, N. PICKARD6, J. GARLAND6 & D. HUNT1 1

Statoil Research Centre, Sandsliveien 90, N5020, Bergen, Norway 2

National Iranian Oil Company, Jomhouri Ave, Tehran, Iran

3

Shell International Exploration and Production B.V., The Netherlands 4

OMV Exploration & Production, Vienna, Austria

5

Institute of Earth Sciences, CSIC, Barcelona, Spain

6

Cambridge Carbonates Ltd, Northampton House, Solihull, UK *Corresponding author (e-mail: [email protected])

Abstract: The Barremian–Aptian upper Khami Group and Albian–Campanian Bangestan Group have been studied at outcrop in Lurestan, SW Iran. The upper Khami Group comprises a thin deltaic wedge (Gadvan Fm) transgressively overlain by shelfal carbonates (Dariyan Fm). The Dariyan Fm can be divided into lower and upper units separated by a major intra-Aptian fracture-controlled karst. The top of the Daryian Fm is capped by the Arabian plate-wide Aptian– Albian unconformity. The overlying Bangestan Group includes the Kazhdumi, Sarvak, Surgah and Ilam formations. The Kazhdumi Fm represents a mixed carbonate-clastic intrashelf basin succession, and passes laterally (towards the NE) into a low-angle Orbitolina-dominated muddy carbonate ramp/shoal (Mauddud Mbr). The Mauddud Mbr is capped by an angular unconformity and karst of latest Albian–earliest Cenomanian age. The overlying Sarvak Fm comprises both lowangle ramp and steeper dipping (5 –108) carbonate shelf/platform systems. Three regionally extensive karst surfaces are developed in the latest Cenomanian –Turonian interval of the Sarvak Fm, and are interpreted to be related to flexure of the Arabian plate margin due to the initiation of intra-oceanic deformation. The Surgah and Ilam Fm represent clastic and muddy carbonate ramp depositional systems respectively. Both The Khami and Bangestan groups have been affected by spectacularly exposed fracturecontrolled dolomitization. Dolomite bodies are 100 m to several km in width, have plume-like geometry, with both fracture (fault/joint) and gradational diagenetic contacts with undolomitized country rock. Sheets of dolomite extend away from dolomite bodies along steeply dipping fault/ joint zones, and as strata-bound bodies preferentially following specific depositional/diagenetic facies or stratal surfaces. There is a close link between primary depositional architecture/facies and secondary dolomitization. Vertical barriers to dolomitization are low permeability mudstones, below which dolomitizing fluids moved laterally. Where these barriers are cut by faults and fracture corridors, dolomitization can be observed to have advanced upwards, indicating that faults and joints were fluid migration conduits. Comparisons to Jurassic– Cenozoic dolomites elsewhere in Iran, Palaeozoic dolomites of North America and Neogene dolomites of the Gulf of Suez indicate striking textural, paragenetic and outcrop-scale similarities. These data imply a common fracture-controlled dolomitization process is applicable regardless of tectonic setting (compressional, transtensional and extensional).

The role and importance of fractures during dolomitization has been the focus of discussion within both academia and industry in recent years (See review in Davies & Smith 2006; Roure et al. 2005; Machel & Lonnee 2002;

Machel 2004). Debate has focused on the following key areas: (a) Do fractures act as conduits for the flow of dolomitizing fluids?

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 343–396. DOI: 10.1144/SP329.14 0305-8719/10/$15.00 # The Geological Society of London 2010.

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(b) Are fractures capable of introducing hotter (hydrothermal) fluids from depth into the surrounding country rock, resulting in the development of regional scale dolomite bodies? (c) From a reservoir perspective: what is the form and vertical-horizontal length scale distribution of the dolomite bodies, when is porosity created and preserved; how is porosity spatially distributed within the dolomite bodies; how does dolomite distribution and porosity relate to primary facies and depositional architecture; and what is the relationship between the dolomite bodies and ‘leached’ microporous limestone. To date, the majority of published data aimed at answering these questions has focused on the integration of seismic, core and outcrop data from the Palaeozoic of North America, with an emphasis on diagenetic/petrographic/geochemical studies to produce conceptual models of dolomitization. The results of these studies have been published as a set of thematic papers (Bulletin AAPG, 90, 2006, see review article by Smith & Davies 2006). However, despite the growing volume of work on fracturecontrolled dolomites, relatively little has been published describing well exposed field examples where detailed cross-cutting field relationships can be systematically walked/mapped out and sampled to give a pseudo-3D picture of a major dolomite body integrating primary facies, sequence architecture, fracture development, diagenetic phases and dolomite body size/orientation. For the reservoir geologist faced with fracture-controlled dolomite reservoirs, this lack of a well documented outcrop case study results in significant uncertainty when trying to populate reservoir models between the core scale and seismic scale, particularly during early field development/appraisal stages (e.g. Grammer et al. 2004; Sharp et al. 2006). In addition, with the exception of the Jurassic Arab, Sargalu and Surmeh formations (Cantrell et al. 2004; Goff 2005), and recently described outcrop examples from Borneo and Northern Spain (Wilson et al. 2007; Lapponi et al. submitted; Rosales & PerezGarcia 2010), the majority of case studies for fracture-fed dolomitization have addressed the Palaeozoic of the North American Craton. The aim of this paper is thus two-fold: (a) To document and describe the facies, depositional and sequence stratigraphic architecture of the ‘host’ Upper Khami and Bangestan groups in Lurestan province, SW Iran. This framework forms the basis for understanding the distribution of dolomite. To date, few published studies have addressed these groups (e.g. James & Wynd 1965; Setudehnia 1978; Taghavi et al. 2006, 2007; van Buchem et al.

2006; Razin et al. 2010). A modern facies and sequence stratigraphic framework was thus required to serve as a reference model for these hydrocarbon prolific reservoirs units in the Lurestan region of SW Iran. (b) To describe and illustrate exceptionally well exposed fracture-controlled dolomite bodies, with specific focus on description and interpretation of field observations and cross cutting relationships between dolomites, fractures, stratal architecture and primary and secondary facies. Although emphasizing the field relationships, our outcrop observations of dolomites are supported by detailed petrographic studies (CL, isotopes, fluid inclusions), a summary of which is included here. The main drive for our outcrop work has been to establish a robust reservoir framework for the Cretaceous Khami and Bangestan groups incorporating matrix (depositional, diagenetic) and fracture (joint, fault) heterogeneity that could be used as a predictive tool in the nearby subsurface. The outcrops studied are located less than 20 km away from the prospective Mesopotamian foreland basin (Figs 1 & 2), and expose an almost identical stratigraphic and depositional succession. In addition, the outcrop diagenetic and structural template is directly comparable to the subsurface. Subsurface data are limited to a few wells and spaced 2D seismic. The outcrops thus represent a unique dataset that can be used to reduce uncertainty in subsurface geological models and understanding, both during exploration and field development (Sharp et al. 2006).

Regional setting and methodology The case study focuses on facies and dolomite bodies developed within the Barremian–Aptian Khami Group and the Albian –Campanian Bangestan Group which, after the Miocene aged Asmari Formation, form the most prolific hydrocarbon reservoir units in Iran (Fig. 3, James & Wynd 1965; Hull & Warman 1970; Setudehnia 1978; Beydoun et al. 1992; Alsharhan & Nairn 1997). The outcrops studied are located in the Anaran Anticline, which forms the south-westernmost anticline of the Simply Folded Belt of the Zagros Mountains (Fig. 1, Emami et al. 2010). In this region rapid recent uplift associated with up to 1 km of erosional incision by rivers (Homke et al. 2004; Verge´s 2007; Verge´s et al. 2010; Emami et al. 2010) has resulted in spectacular pseudo-3D outcrops where individual dolomite bodies can be walked out and sampled from core to tip in dip, strike, plan and vertical sections (Figs 2, 4 & 5). In addition, the relationship to primary depositional

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Fig. 1. (a) Geological map of Anaran Anticline (Modified from Emami et al. 2010). The Anaran Anticline forms the western-most fold of the Simply Folded Belt of the Zagros Mountains bounded to the west by the Mesopotamian Foreland Basin. The boxed area in red forms the focus of dolomite bodies studied in this paper (NW Dome) and is enlarged in Figure 2. Dolomites have also been studied at Kuh-E-Pashmi. Location of cross section A–A0 shown. Insert at top right shows regional location of the Anaran Anticline. (b) Simplified cross-section through the central portion of the Anaran Anticline. Bedding dips are indicated. MFF, Mountain Front Flexure.

346 I. SHARP ET AL. Fig. 2. (a) Oblique satellite and DEM view of the NW Dome area, Anaran Anticline (see Fig. 1 for location). Note asymmetric nature of Anaran Anticline, with long shallow-angle dipping NE limb and a short, steeper dipping SW limb. Also note crestal normal faulting. (b) Simplified stratigraphic column of exposed stratigraphy in Anaran Anticline. Interval studied indicated. (c) Photo panorama of NW Dome river gorge. Photo corresponds to red boxed area in Figure 2a. Massive dolomites on the right (NE) of the image pass into limestones on the left (SW). Dolomite bodies are outlined by white dashed line. Also note crestal graben defined by normal offset of top Lower Sarvak (MFS-Sa3 Ahmadi Member). Faults offset dolomite bodies. Logged Sarvak section and Lower Sarvak lithostratigraphic units indicated.

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Fig. 3. Simplified stratigraphic column of the Cretaceous in SW Iran. Modified from James & Wynd (1965). This study focuses on outcrops in the Lurestan area.

facies, sequence architecture, diagenetic facies and fractures ( joints and faults, Fig. 6) may be studied and documented in detail. Horizontal (plan section) outcrops of the host limestones and dolomite bodies comprise large bedding plane exposures from which it was possible to capture joint and fault density and dolomite distribution by analysis of both high resolution satellite data (IKONOS, QuickBird) and geo-rectified high resolution digital photographs taken from a helicopter (Figs 5 & 7a). Multiple cross-sectional outcrops

are afforded by antecedent river systems and associated tributary drainage which have cut through the core of the anticline, giving excellent full section outcrops in a variety of orientations (i.e. both depositional dip and strike, Figs 2, 4, 5 & 6a). The sections were captured by high resolution digital photography which were geo-located on a satellite draped DEM of the study area. LiDAR data have also been obtained. Vertical stratigraphic logging and sampling was undertaken to establish several control sections

Fig. 4. Helicopter photograph of conjugate normal faults (sub-vertical white dashed lines) and spectacularly exposed dolomites bodies (edges outlined by white dashed lines), SW flank of Anaran Anticline, NW Dome. View is towards NW. Massive dolomites on the right (NE) pass into stratabound dolomites and limestones on the left (SW). Note rotated (folded) aspect of normal faults (rotation towards left/SW), and that dolomite bodies are offset by the faults. Exposed section is 750 m thick. A hanging wall syncline is developed in the Ilam Fm along the SW fault due to ductile folding of the Surgah Fm (see also Fig. 16a).

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Fig. 5. (a & b) QuickBird images, NW Dome, river gorge section (see also Fig. 2c). Dolomite bodies visible in dark brown. Trace of NW–SE trending normal faults which offset the dolomite bodies indicated by red arrows. Location of Figures 5b, 7a & 18 indicated. View points for Figures 2 & 4 also indicated. Red-boxed area indicates Figure 5b location. Note isolated dolomite plumes developed in the Upper Sarvak Fm in right part of image. (b) Detail of Figure 5a, Contact between massive dolomites (right) to stratabound dolomites and limestones (left) indicated by white dotted line. Top lower Sarvak/MFS-Sa3 (Ahmadi Mbr) marked in blue. The Ahmadi Mbr is interpreted to have formed a significant barrier to vertical migration of dolomitizing fluids. This contact is offset by two normal faults (red arrows). See Figure 4 for outcrop view of same area. Location of Figures 20a & 23 indicated.

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Fig. 6. (a) Georectified outcrop photograph of Lower Sarvak Fm (sequences Sa1 and Sa2, Fig. 10). Strike of face is NE-SW, thus NW striking joint trend is well represented. Joints and small faults (red lines) and major bedding breaks (yellow) have been digitized. Note variation in joint density from lithology to lithology (high joint density in rudistic grainstones of top LSF. Low joint density in lagoonal wackestones of LSE). Where a series small faults (0.5–2 m displacement) cut the section, a through going damage zone is developed in all lithologies, resulting in good vertical fracture communication. Note stratabound D2 dolomite in unit LSC (dark brown colour). (b) Top Mauddud Mbr (unit H) and Lower Sarvak Fm (Units A– G) joint density plots by lithostratigraphic unit, NW Dome type section. Density data based on outcrop line sampling and digitized outcrop photographs. Highest density (upto 14 joints per metre) are within the more ‘grainy’ lithologies (LSF platform margin rudistic shoal, LSE2-1 platform top grainstone tidal shoal, LSD 2-2 and 1-1 rudist-foram platform margin shoal, LSH Mauddud Mbr karst). Lowest joint density (less than 1 joint per metre) is within platform top/lagoonal micrites of units LSE 1-1, 2-3, 3-2, 3-4, 3-6.

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Fig. 7. (a) Georectified helicopter photograph looking vertically down onto joints in basinal-slope wackestones of Upper Sarvak Fm (sequence Sa6), NW Dome. Jointing is predominantly NE and NW striking. Density of NE set is twice density of NW set. Joint density is ‘low’ by comparison to packstone, grainstone and dolomite lithologies. NW– SE striking set abut against NE– SW striking joint set, implying that the NE– SW set formed first. (b) Simplified representation of jointing developed in Sarvak Fm. Jointing is developed at 908 to bedding, and joint density varies markedly from layer to layer. Joints stop at bedding planes, particularly at depositional cycle tops or marly/claystone layers. Joints are thus predominantly stratabound, but occasionally go through more than one layer. Joint density in NE orientation is twice that in NW orientation. (c) Outcrop reservoir model (350 m high, 1.2  1.6 km wide) of Lower Sarvak Fm in NW Dome. Joint density is shown. Note vertical joint heterogeneity and horizontal joint homogeneity. See Sharp et al. (2006) for a full description.

from which unit thickness, facies, microfacies, poretype, porosity–permeability data, joint density data per facies type and reservoir zonation, could be derived. Porosity-permeability typing was based on thin section study using the methodology of Lønøy (2006) and conventional analysis of plugs drilled at outcrop or collected from large samples. ‘Walking-out’ and sampling of depositional and diagenetic elements away from the control sections captured horizontal and vertical correlation length scales of primary and secondary facies and, in part, porosity-permeability distribution. Similarly, ‘walking out’ and digital photo line drawing interpretation captured depositional architecture, diagenetic and fracture heterogeneity (e.g. Fig. 6a). An emphasis was placed on the systematic documentation of paragenetic and textural relationships at all studied outcrops, in particular the relationship between dolomitized units and fractures ( joints and faults). Petrographic work has involved standard thin section analysis, cathodoluminesence and scanning electron microscopy. Geochemical work has involved stable isotope analysis (d13C, d18O,

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Sr/86Sr ratio, trace elements) and fluid inclusion microthermometry. This work has allowed isolation of individual cement phases and development of a fluid and cement stratigraphy, a summary of which is included here. The final stage of the study has involved building an outcrop reservoir model. The work flow and methodology used during reservoir model building is documented in Sharp et al. (2006).

Structural framework The Anaran Anticline is a 100 km-long and 5 km wide mountain-forming anticline that forms the frontal fold of the Zagros Fold Belt in the Lurestan region of SW Iran (Fig. 1). The anticline has a maximum topography of about 1.6 km in its central segment. To the SW of the Anaran Anticline lies the hydrocarbon-prolific Mesopotamian Foreland Basin of Iran and Iraq, and to the NE lies the spectacularly exposed Zagros Simply Folded Zone (Blanc et al. 2003; McQuarrie 2004; Sepehr & Cosgrove 2004; Sepehr et al. 2006; Sherkati &

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Letouzey 2004; Sherkati et al. 2005; Verge´s et al. 2010). The elevation difference between the Khami and Bangestan groups exposed at outcrop and in the subsurface is between 4– 5 km, achieved along a major basement-involved steep normal to reverse fault and/or monoclinal flexure, the Mountain Front Fault/Flexure (Figs 1b & 8, Blanc et al. 2003; Emami et al. 2010; Verge´s et al. 2009). There are pronounced changes in structural orientation along the strike of both the Anaran Anticline and anticlines imaged on 2D seismic data in the foreland (Fig. 1a). These changes in orientation are thought to be related to a combination of underling N –S and NE–SW trending basement faults, and/or interference and linkage of originally isolated fold segments during fold growth and fault propagation (Emami et al. 2010). Fold development in both the prospective foreland and the studied outcrops is related to the collision of the Eurasian plate and the Arabian plate in the Miocene. In the study area, fold development is constrained to have initiated after 8 Ma, as documented by growth strata in the Changuleh and Zarrin Abad synclines which onlap the Anaran Anticline (Fig. 1, Homke et al. 2004). Hydrocarbon generation and migration occurred slightly prior to and coeval with folding (Bordenave & Hegre 2005). In both the outcrops and the proximal foreland, the structural style is of open asymmetric folds with a vergence towards the SW (towards the foreland). The NE limbs typically dip between 58 and 158 at Bangestan levels, whilst the SW limbs dip between 158 and 368 (Fig. 1b). The dolomite outcrops described in detail in this paper are located at the northern end of the Anaran Anticline (NW Dome, Kuh-E-Pashmi, NS Fault Complex; Figs 1 & 2).

Fig. 8. 3D sketch of Anaran Anticline combining outcrop and seismic data to illustrate main structural features. Section is representative of central portion of Anaran Anticline (Fig. 1). Note NW –SE trending F1 conjugate normal fault set, which are passively rotated (folded) and offset by NE– SW trending F2 faults.

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During our study we made a clear distinction between faults and joints, which collectively define fractures. The reason for this is three fold: (a) Faults are through-going structures associated with significant damage zones and displacement which offset stratigraphic/reservoir layers. They represent major conduits for flow, and thus may be expected to have had a major impact on the migration of diagenetic and hydrocarbon-bearing fluids. (b) Joints are short dilational fractures upon which there is no visible shear displacement. As such they are not major flow conduits, but represent discontinuities along which both diagenetic and hydrocarbon-bearing fluids could access volumetrically large areas of undolomitized country rock. (c) The timing of development of joints and faults is not necessarily synchronous, and thus they could record different diagenetic histories. Despite the overall contractional setting, the predominant small-scale structures observed in the Khami and Bangestan groups at outcrop are normal faults (Figs 1, 2 & 4). Bedding parallel slip is rare or absent, with the exception of the clay-rich Surgah Fm and limestone-marl Ilam Fm (Fig. 4). The Surgah Fm forms an excellent smear horizon along faults. Fault displacements are typically in the order of 50 –300 m, although throws of upto 1000 m have been recorded (N –S Fault Complex, Fig. 1a). Displacement on normal faults within the Sarvak and older formations is often lost up-section due to accommodation of displacement by folding and bedding parallel slip within the overlying Surgah and Ilam fm’s. The Surgah Fm thus forms an excellent regional top seal to the vertical migration of hydrocarbon and dolomite/Mg-bearing fluids. Remote sensing and field mapping indicate that there is a greater concentration of normal faulting on the steeper dipping SW limb, and that fault development was slightly pre- to coeval with folding (Figs 1, 2, 4 & 5). 2D seismic data in the foreland indicate a comparable development of normal faults in relation to folding. Fault kinematics are complex, and at any given location two sets of conjugate normal faults can be recognized (NW– SE: F1, NE –SW: F2, Fig. 8) with two extension directions which vary in orientation along the anticline. Pre-last phase of folding, NW– SE striking, conjugate faults can be recognized, as they have been passively rotated towards the foreland, with SW dipping faults showing increased dips (locally overturned) and NE dipping faults showing decreased dips (Figs 4, 8 & 9). The geometries are similar to rotated normal faults described from extensional settings by Sharp et al. (2000). All studied normal faults are associated with significant

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Fig. 9. Conceptual structural evolution of the Anaran Anticline. North is towards top of page.

damage zones and fault rock, and have significantly increased permeability in relation to the surrounding country rock. The majority of observed fault planes are associated with vertical to slightly (58) oblique slickensides and corrugations, although rare pure oblique (strike) slip faults have been observed. Joints are pervasive at outcrop (Figs 6 & 7), and are typically stratabound or partially stratabound. In map view, the joints form well connected networks (Fig. 7). Two joint sets predominate, which strike NE–SW and NW–SE. The NW–SE set often abut against the NE –SW set, indicating that the NE –SW set pre-dates the NW –SE set. Earthquake focal mechanisms in the region (Hessami et al. 2006) indicate that the principal horizontal stress is NE –SW. This orientation is thus the optimal direction for open fractures. Borehole break-out and image log data from wells in the foreland indicate a comparable joint and in-situ stress pattern. Joint density within any one lithologically homogenous stratigraphic unit is relatively uniform across the Anaran Anticline. However, joint density is very heterogeneous vertically showing an organized and predictable lithology/facies control on spacing and density (Figs 6 & 7). Joint

density is typically highest in grainstones and dolomites and is lower in carbonate mudstones. Shales or marls are unjointed or sparsely jointed and tend to form vertical barriers to joint propagation. Such shales and marls could thus also have formed significant barriers/baffles to fluid flow (dolomitizing fluids and hydrocarbons). Joints typically are developed perpendicular to bedding, suggesting they formed prior to significant folding. Also, in most of the region, faults appear to have no effect on joint orientation, implying that the joints pre-dated the faults. However, in one well exposed area (Ghir-ab, Fig. 1a), the first, most continuous joint set changes orientation systematically by over 458 in the vicinity of a fault, so that close to the fault, the joints and the faults have the same strike. This indicates that the joints formed during or after faulting in this area (Rawnsley et al. 1992). Whilst there is clearly a continuum between jointing, faulting and folding, the majority of our outcrop data indicate the following deformation sequence (Fig. 9): (a) First phase of joint formation was prior to both significant folding and faulting. Joints form at 908 to bedding.

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(b) Initial phase of faulting was pre- to syn-folding associated with layer parallel extension. (c) The last phase of folding passively rotated (folded) NW–SE trending faults, and was associated with NE–SW trending faulting which offset NW–SE trending faults. This last phase of faulting was also associated with fold tightening and uplift above the Mountain Front Flexure (Emami et al. 2010).

Stratigraphic framework The Barremian–Aptian Khami Group and Albian – Campanian Bangestan Group were deposited on the NE facing passive margin of the Arabian Plate and are dominated by carbonate lithologies with subordinate interbedded clastics (distal deltaic and shelfal shales, silts and sands, Fig. 3). A recurring lateral facies association of clastics – intrashelf basin – carbonate shelf – open marine basin characterizes both groups moving from the Arabian Shield in the SW (Iraq) towards the open marine Tethys in the NE (Iran), Davies et al. (2002). Since the publications of James & Wynd (1965) and Setudehnia (1978), little has been published on the sedimentological and stratigraphic development of the Khami and Bangestan groups in Iran. Correlation of the Iranian units within an Arabian plate sequence stratigraphic framework was attempted by Sharland et al. (2001), and van Buchem et al. (2006) have proposed a correlation scheme for the Cretaceous of SW Iran. Razin et al. (this volume) and Taghavi et al. (2006, 2007) have recently suggested a revised stratigraphic and reservoir framework for the Sarvak Fm based on study of outcrops in the High Zagros and core material from the Dehluran field. Detailed biostratigraphical, sedimentological and sequence stratigraphical studies undertaken during the course of this study are summarized in Figure 10 and in the following section on a formation by formation basis (e.g. Fig. 11). The framework is based on integration of outcrop and subsurface data throughout Lurestan, although in this paper emphasis is placed on description of stratigraphic development in the Anaran Anticline. This framework forms the basis for understanding the distribution of secondary dolomite bodies. To establish the stratigraphical framework emphasis was placed on identification and dating of Maximum Flooding Surfaces (MFSs) as these have the greatest geographical extent, are often datable at outcrop and in the subsurface (using ammonites, nannofossils, palynomorphs and microfossils) and are distinct on wireline logs. However, we define depositional sequences (Sensu Van Wagoner et al. 1990) as opposed to genetic sequences (sensu Galloway 1989), as depositional sequences allow a better tie to

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the existing lithostratigraphic framework (Fig. 10). System tract assignment is based on geometrical observations from seismic-scale outcrops, and inferred geometrical relationships based on correlation of outcrop and well data.

Gadvan Formation The Gadvan Formation forms the deepest stratigraphic exposures of the Khami Group in the Anaran Anticline, and comprises a 25 m thick interval (incomplete section, base not exposed) overlain by massive limestones and dolomites of the Dariyan Formation (Fig. 10). Two distinct facies associations are present; (i) A basal succession of silty-marly nodular bedded Orbitolina sp., and coral-algalrhodolith bearing carbonates and calcarenites. Individual cycles are 1.5 to 3 m thick, comprising nodular bedded marly Orbitolina sp., algae wackestones to packstones overlain by increasingly calcarentic and sandy facies associated with dark organic rich shales. Thalassinoides sp., rhodoliths, iron staining and bored shell debris mark the change from the carbonate to clastic sediments, and are interpreted as firmgrounds and hardgrounds. Some of the overlying calcarenites comprise rounded and abraded limestone clasts, lithologically identical to underlying facies. This indicates erosion and redeposition of the adjacent and/or underlying carbonate succession. An Early Aptian microfauna (KM14 Biozone) was dated from the lower part of the logged interval. Micro and macro fauna present include Palorbitolina spp., Palorbitolina lenticularis, Salpingoporella dinarica, Pseudocyclammina lituus, Lithocodium aggregatum, miliolids, agglutinating foraminifera and common echinoderm debris. (ii) An overlying, increasingly clastic dominated interval of dark laminated shales (fissile, organic-rich, TOC up to 5.8%) and very fine grained highly bioturbated calcareous silty sandstones. The tops of cycles are abrupt and the base of the next cycle is marked by dark organic-rich laminated shales. These cycles are best interpreted as distal deltaic or distal lower shoreface parasequences. Simplistically, an upward transition from a restricted shallow ramp-platform interior/lagoon to an increasingly more ‘proximal’ mixed carbonate-clastic setting appears likely. The clastic facies are interpreted to reflect progradation of a deltaic system (Upper Zubair equivalent) across the ramp interior (i.e. ‘clastic drowning’ of the

Fig. 10. Stratigraphic and reservoir framework for the Upper Khami and Bangestan groups in Lurestan. The scheme is based on logged sections, biostratigrahy (micro and macro fauna), diagenetic, geometric and sequence stratigraphic analysis in the Anaran and adjacent areas (Kabir Kuh, Siah Kuh, Chenareh, Rit, Khorrama-Abad, Anjir and Sultan Anticlines). Where possible, the sequence stratigraphic scheme has been correlated to Sharland et al. (2001) and van Buchem et al. (2006). Inferred periods of active tectonics are indicated, based on outcrop observation of faulting, jointing, tilting and fracture-controlled karstification. The lithostratigraphic units allow correlation on a local scale in wells and outcrops. Joint density is based on facies, and does not take into account location on a fold (e.g. crest vs flank). D1 dolomites are typically developed as dolomitic limestones (scattered dolomite rhombs in limestone matrix), and are commonly associated with stylolites in slope and platform top facies. D2-D3 dolomites are associated with fracture and fold development, and are prevalent beneath aquitards, within karsts, nucleated on earlier D1 dolomite, and within permeable HST strata. The most favourable primary reservoir facies are skeletal grainstones developed during late HST. Secondary reservoir facies develop due to fracturing and late diagenesis. Reservoir quality is related to primary facies (e.g. fractured HST skeletal grainstone margins will typically have better reservoir properties than fractured and dolomitized basin-slope facies). Early cements typically occur within TST packstones and grainstones (marine cements), and have a detrimental effect on reservoir quality. Early cemented intervals are not strongly affected by later diagenesis (dolomitization). Silica cements are pore filling or developed as chert nodules, and have a detrimental effect on reservoir quality. Meteoric diagenesis is associated with HST-FRST-LST and karsts. Both reservoir enhancing and destroying affects occur. Source Rocks are of two depositional types: (i) Pro-delta plant-rich mudstones (e.g. Gadvan, Nahr Umr); (ii) Restricted intrashelf basin facies and MFS intervals (e.g. Ahmadi and Ghirab Mbrs).

Fig. 11. Logged stratigraphic sections of the Dariyan, Kazhdumi, and Lower Sarvak formations in NW Dome and N-S Fault Complex. See Figure 1 for location. Main depositional sequences and facies indicated.

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carbonate environment). This interpretation is in agreement with the regional depositional models of Sharland et al. (2001) and Davies et al. (2002). In our sequence stratigraphic scheme a sequence boundary (SB-Ga1 – candidate K70 SB of Sharland et al. 2001) is placed near the top of the Gadvan Fm within the maximum regressive clastic deposits.

Dariyan Formation The Dariyan Fm is 163 m thick where fully exposed in the NW Dome section and totally dolomitized. Early Aptian ages have been confirmed from an undolomitized section further to the SE (N –S Fault Complex, Fig. 11). The Dariyan Fm can be divided into Lower and Upper, separated by a major karst surface, SB-Da1 (Figs 10 –12). An angular unconformity and karst surface also caps the Dariyan Fm, marking the Aptian –Albian boundary and the contact between the Khami and Bangestan groups, SB-Da2 (Fig. 12). The Late Aptian has not been positively identified, although

it is speculated that the Upper Dariyan Fm is at least partly Late Aptian in age. The Lower Dariyan Fm is 100 m thick and abruptly overlies the Gadvan Formation. This contact is interpreted as a major transgression (MFS-Da1, candidate Earliest Aptian MFS K70 of Sharland et al. 2001, Fig. 10). Despite dolomitization primary facies may be identified. The Lower Dariyan Fm comprises prograding grainstone shoal units overlying muddy Orbitolina sp., and coral-algal bearing nodular wackestones. Individual shoal cycles comprise nodular bedded bottom/toesets, and distinctly cross-bedded forsets to topsets. Shoals can be up to 15 m thick, and are typically rich in reworked rudistic material (Agrioplurids sp.,) in their upper part. Bioturbated cycle tops are picked out by Thalassinoides sp. Overlying facies comprise nodular bedded wackestones with intensive bioturbation and miliolids evident in thin section. A low angle tidally influenced ramp margin to inner ramp setting is envisaged. In the undolomitized N– S Fault Complex

Fig. 12. Outcrop photograph of angular contact between the Khami and Bangestan groups, NW Dome gorge section. Faults and joints are karstified beneath SB-Da1 and SB-Da2, indicating at least two periods of active tectonics in the upper part of the Khami Group. Section is partly dolomitized. Location of NW Dome section indicated on left. The lower 2 lithostratigraphic units of the Kazhdumi Fm are also well exposed. Exposed total section is c. 150 m thick.

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section (Fig. 11) an inner ramp/platform top (lagoonal) depositional setting is confirmed based on the occurrence of abundant miliolids, Palorbotolina lenticularis and Salpingoporella dinarica. The top of the Lower Dariyan Fm is marked by a major fracture-controlled karst (SB-Da1, Figs 10–12). In the NW Dome section solutionenlarged E–W and NW–SE trending karstified joints (1–8 m wide) can be mapped extending up to 50 m into the underlying Lower Dariyan Fm. The fractures are filled by intraformational breccias, vadose silts and pebble conglomerates. Small stratiform karstic caverns, up to 1 m high, can be mapped away from the main fracture zones. Individual caverns have very irregular solution controlled bases and a composite fill, comprising laminated vadose and locally pisolithic cave sediments at the base overlain by radial sparry cements. These karstic features are totally dolomitized. The top 10–15 m of the Lower Dariyan Fm is characterized by a rubbly weathering karstic breccia, although later overprint by D2 and D3 –D5 dolomites (discussed later) obscures much of the primary karst fabrics. The Upper Dariyan Fm is between 35–65 m thick (Fig. 11). The base is marked by a transgressive, 8 m thick package of iron-stained, glauconitic, bioturbated, oyster and rudist-rich tabular bedded packstones and grainstones with bored intraclasts which lie horizontally on the dipping and rubbly weathering karst zone below. These beds locally ‘sag’ downwards into the underlying karst zone. In the overlying interval, large cross-bedded rudistfilled sub-tidal channels (upto 20 m wide) are developed, overlain by an homogenous interval rich in Orbitolina-bearing wacke-packstones with bioturbated bed tops (Thalassinoides sp.,). The upper 11 m of the Dariyan Fm is intensly bioturbated and rich in small conical Orbitolina sp., capped abruptly by a marl/clay layer. This contact marks the top of the Dariyan Formation, and is associated with a significant angular uncomformity (between 1–58 dip discordance between Dariyan Fm and overlying Kazhdumi Fm, SB-Da2 – Fig. 12). Biostratigraphic dating of the N –S Fault Complex section indicates that the topmost Dariyan Fm is of Early Aptian age (KM14 biozone) and is overlain by a thin Early Albian interval at the base of the Kazhdumi Fm. The Late Aptian has not been positively identified. The presence of an angular unconformity plus conjugate joints (with minor karst) at the top of the Dariyan Fm is interpreted to indicate a period of normal faulting. This unconformity is well known on the Arabian Plate (Sharland et al. 2001).

Kazhdumi Formation and Mauddud Member The Kazhdumi Formation is 50 m thick at outcrop in the Anaran Anticline and up to 150 m thick in

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nearby wells to the SW. It is entirely Albian in age, and typically comprises a regionally developed three-fold lithostratigraphic succession (Figs 10 & 11): (i) A lower recessive weathering Orbitolina limestone marl-dominated interval overlain by glauconitic calcarenites – Unit KZC (ii) A central unit of Orbitolina and pelloid-rich grainstones – Unit KZB (iii) An upper recessive weathering interval of organic-rich calcareous mudstones – Unit KZA. A shallow water (lagoonal to mid ramp) setting is suggested for unit KZC, with abundant conical Orbitolina sp., possibly indicative of a sea grass meadow setting (cf. Davies 1970; van Buchem et al. 2000a, 2002a; Simmons et al. 2000; Jones et al. 2004). The first few metres of this interval, however, yield belemnite guards, pectinids and large discoidal Orbitolina sp., representing a major flooding event over the top Dariyan unconformity. Biostratigraphical data indicate an Early Albian age, although the first sample above the top Dariyan unconformity is of possible Late Aptian age (?KM13 biozone) in the N –S Fault Complex section. Towards the upper part of unit KZC, two regionally developed 1 m thick Orbitolina-bearing micrite-rich layers are capped by an iron-phosphate stained surface. Overlying marls (3 m thick) are rich in planktonic dinoflagellate cysts and planktonic foraminifera, and show a decline in orbitolinids and miliolids. This interval is picked as an MFS (MFS-Kz1, Early Albian K90 MFS candidate of Sharland et al. 2001, Figs 10 & 11). The MFS is abruptly overlain by silt and very-fine grained sandstone rich in glauconite and conical Orbitolina sp. Well exposed tidal bundling within tidal bars with robust Ophiomorpha nodosa extending off set boundaries are evident. Conical Orbitolina sp. in this unit are abraded and reworked. A very shallow water depositional setting (subtidal, proximal ramp interior) appears likely, and the contact with the marls below is regressive. The sands are overlain by a 15 m thick resistant weathering cross-bedded grainstone shoal complex (Unit KZB) rich in pelloids, ooids, red algae and abraded Orbitolina sp. (Fig. 11). Angular clasts derived from the underlying unit are also present, associated with a fungal spore peak, and may be interpreted as a regressive event. SB-Kz1 may be picked at this fungal spore peak or at the base of the glauconitic sands below (equivalent to regressive maximum of Burgan/Nahr Umr Deltaics, Fig. 10). Above this, the limestones ‘clean’ upwards and are overlain by a thin interval of marly recessive-weathering limestones extremely rich in conical Orbitolina sp., high spiral gastropods and

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in-faunal echinoderms. This interval is capped by an iron oolite bed, which marks a distinct deepening of the depositional environment (? Transgressive Surface) associated with an increase in dinoflagellates and decrease in fungi. The upper lithostratigraphic interval of the Kazhdumi Formation (unit KZA, 20–30 m thick at outcrop, 50 m thick in subsurface) is recessive weathering and essentially records the ‘deepest water’ facies of the Kazhdumi Fm. The basal iron oolite is interpreted as a hard-ground, and is overlain by pectinids and Orbitolina sp. bearing marls and a second hard-ground associated with Lithophaga sp., bored iron carbonate clasts. The uppermost units comprise organic-rich, thin bedded micritic limestones and marls which are very rich in dinoflagellates and in-faunal spiny echinoderms. This upper interval is interpreted as the most ‘distal’ facies (intrashelf basin/outer ramp). Biostratigraphically, the interval above the iron oolite is of Middle Albian age. A MFS is picked just above the second hardground associated with maximum marine signature based on relative increases in planktonic foraminifera, dinoflagellate cysts and foraminiferal test linings (MFS-Kz2, candidate Middle Albian MFS K100 of Sharland et al. 2001, Fig. 10). The lithostratigraphical contact between typical Kazhdumi Fm facies of unit KZA and the base of the Mauddud Mbr is marked by a rapid shallowing to cliff-forming nodular bedded and cross-bedded tidal shoals (Fig. 11). A thin mudstone interval just below this contact is dated as Late Albian in age, associated with a relative increase in planktonic foraminifera, dinoflagellate cyst abundance, foraminiferal test linings and decline in orbitolinids. This localized event may be part of the underlying deeper water interval seen within the Middle Albian section. However, it may represent a separate phase of increased water depth and as such is a possible candidate level for placement of the K110 MFS of Sharland et al. (2001). This is MFS-Kz3 (Fig. 10). At outcrop, MFS-Kz3 is abruptly overlain by shallow water facies of the Mauddud Mbr, whilst in nearby wells to the S/SW a thick organic-rich Late Aptian succession is present, characterized by relatively common Favusella washitehsis. This facies configuration is similar to that proposed by Bordenave & Burwood (1995) and Bordenave & Hegre (2005), with a shallow water Mauddud Mbr carbonate shoal (Bala Rud Shoal) in the area of the Anaran Anticline –Bala Rud Flexure separating the Garau Basin to the north from the Kazhdumi intrashelf basin to the south. Progradation of the Mauddud Mbr appears to have been both towards the open Garau Basin in the NE and the Kazhdumi intrashelf basin in the SW. At outcrop in the Anaran Anticline, the Kazhdumi to Mauddud transition represents a relatively

simple prograding highstand succession. Vertically three main facies associations are developed; (a) Thin-bedded organic-rich micritic limestones and marls, interpreted as intrashelf basin/ deep ramp deposits (Kazhdumi Fm unit KZA). (b) Mid ramp facies (Mauddud Mbr). Orbitolina sp., and echinoderm-rich wackestones, packstones and grainstones deposited as high energy subtidal shoals characterized by nodular bedded bioturbated bottomsets and bi-directional tabular to trough cross-bedded foresets to topsets. Thalassinoides sp. and Ophiomorpha sp. burrows are common extending downwards off bed and set boundaries. The burrows typically are dolomitized. Individual shoal cycles in this facies association are 5–10 m thick. (c) Inner ramp (Mauddud Mbr). Monotonous Orbitolina sp. wackestones to packstones associated with locally developed layers rich in elevator rudists, Chondrodonta sp., and rare solitary corals. Rudists and Chondrodonta sp., typically are reworked as storm lags, although they locally occur in-situ. Dolomitized Thalassinoides sp. burrows are very common, and typically extend off bed boundaries. In-faunal echinoderms are also present. Secondary replacement chert and dolomite are relatively common in the inner ramp facies. A major karst (SB-Kz3) caps the Mauddud Mbr and has a penetration depth/profile of 15– 20 m (Fig. 13). The karst is fracture-controlled by early NW– SE trending joints and normal faults. These joints and faults are transgressively ‘sealed’ by the overlying Sarvak Formation. Erosional relief along the top Mauddud Mbr karst surface is in the order of 3– 9 m, with a low amplitude sink hole/dolinelike morphology mappable in well exposed sections. Sink holes and dolines show a preferential elongation parallel to NW–SE trending faults and joints, with a preferential development in the hanging wall of faults (Fig. 13). A sheet-like karstic solution breccia is locally developed below the karst surface, and varies from 5 m to 10 cm in thickness. Replacement chert and dolomite commonly fill secondary porosity within the karst breccia (Fig. 13). A reddish mudstone is also evident above the breccia, and forms a locally important permeability barrier. Below this uppermost heterogeneous breccia level, the karst is typically expressed as solution-enlarged fractures which pass into two levels of stratabound caverns occurring along major bedding breaks. Caverns and fractures are partially filled by reddish vadose silt and clay, which are themselves cut by dolomite filled fractures. The caverns are upto 5 m across and 1–2 m high, typically with flat bases and convex roofs. Vuggy porosity is locally developed between the solution-enlarged fractures

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Fig. 13. Outcrop photo of the contact between the Mauddud Mbr and base Lower Sarvak Fm in NW Dome. The contact is marked by a well developed fracture-controlled karst (red dashed line). Significant karstic relief is evident picked out by sink holes/dolines, which are elongated parallel to and in the hanging wall of NW– SE trending faults (trace of fault indicated on left). Solution enlarged fractures extend down from the karst surface, and link to two major levels of stratabound caves. The karst has been extensively overprinted by D2 and D3– D5 dolomites, which utilized the karst as a major permeability pathway. D2 dolomitization progressively dies out upwards into muddier and less fractured facies of lithostratigraphic unit LSG of the Lower Sarvak Fm.

and caverns, with vugs either developed along hairline fractures or forming enlarged moulds after dissolved bioclasts. Correlation to nearby wells to the south indicates the presence of a detrital unit (up to 50 m thick) which lithostratigraphically sits between the Kazhdumi and Lower Sarvak formations. This detrital unit (lithostratigraphic unit LSH, Fig. 10) is rich in abraded and well rounded clasts of Mauddud Mbr affinities set within basinal facies. The unit is dated as latest Albian-Earliest Cenomanian, and is interpreted as a lowstand wedge (LST) shed into a remnant Kazhdumi intrashelf basin during erosion of the uplifted and karstified Mauddud Mbr high to the NE. Inclusion of the Mauddud Mbr within the Sarvak Fm thus needs revision, as these new data indicate that the Mauddud Mbr is genetically related to the Kazhdumi Fm as a lateral facies equivalent, and separated from the overlying Sarvak Fm by a major karst. This interpretation is in

keeping with observations by van Buchem et al. (2006, pers comm. 2005) from the Fars province of SW Iran.

Sarvak Formation The Sarvak Formation has been informally divided into the Lower and Upper in the Lurestan area (Fig. 3, James & Wynd 1965; Player et al. 1966; Setudehnia 1978). The contact between the Lower and Upper Sarvak is topographically well expressed throughout Lurestan, typically separating massive neritic cliff-forming carbonates of the Lower Sarvak Formation below from more recessive weathering basinal carbonates and mudstones of the Upper Sarvak Formation above, (Figs 2, 4, 14, 15 & 16a). Foraminifer and ammonite fauna collected during the course of this study allow assignment of this boundary to a latest Early Cenomanian flooding event (MFS-Sa3, Ahmadi Member).

Fig. 14. Logged stratigraphic section of the Upper Sarvak, Surgah and basal Ilam formations in NW Dome. See Figure 1 for location. Main depositional sequences and facies indicated. The Upper Sarvak is developed in relatively ‘distal’ basin and slope facies in NW Dome, but passes into more ‘proximal’ higher energy shelf margin and platform interior facies towards the SE in Central and Southern Anaran (e.g. Fig. 16b). SB-Sa4 and sequence Sa5 are poorly developed in the NW Dome section.

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Fig. 15. Outcrop photo of type section of Mauddud Mbr and Lower Sarvak Fm exposed in NW Dome, Anaran Anticline. Depositional sequences and lithostratigraphic units shown. Cliff face is 350 m high. Depositional sequence Sa1 is well expressed in this section, with a strongly progradational signature from slope/outer platform wackestones (unit G), to coral-rudist platform margin shoals (unit F) overlain by a thick platform top succession (unit E). Note the extremely fractured nature of unit F (Rudist shoals). Sequence Sa2 is developed in stacked platform margin to platform top facies capped by 2 micro-karst intervals (top units C and B). SB-Sa2 is placed at the contact between lithostratigraphic units LSB and LSA.

Lower Sarvak Formation The Lower Sarvak Formation transgressively overlies the top Mauddud Mbr karst and can be divided into two depositional sequences; Sarvak Sequence 1 (Sa1) and Sarvak Sequence 2 (Sa2, Figs 10, 11, 14 & 15). Sa1 is of ?latest Albian-earliest Cenomanian age, and Sa2 is entirely of Early Cenomanian age. The lithostratigraphic base of the Lower Sarvak is associated with a relatively thin transgressive system tract (TST, 30 m thick) which infills the erosional relief of the top Mauddud Mbr karst overlain by a thick highstand system tract (HST, 150 m). A lowstand system tract (LST, lithostratigraphical unit LSH) can be identified in the remnant Kazhdumi intrashelf basin to the SW and Garau Basin to the NE, where it is typically rich in secondary silica (chert) and dolomite. The TST comprises two main facies associations (Fig. 11):

(a) Rudist –Chondrodonta shell banks and biostromes, abruptly overlain by (b) outer platform Orbitolina-rich bioturbated wackestones-packstones. MFS conditions (MFS-Sa1) are associated with closely spaced omission surfaces with abundant Thalassinoides bioturbation (dolomitized). In the subsurface, the TST and associated MFS-Sa1 interval is thicker and associated with a rich planktonic faunal association, including relatively common Favusella washithensis. Above MFS-Sa1, the HST is associated with a strongly progradational interval characterized by a three-fold facies association (Figs 10, 11 & 15): (a) Outer platform Orbitolina-rich bioturbated wackestones-packstones (lithostratigraphic unit LSG) (b) Platform margin rudistic and pelloidal-benthic foram shoals (lithostratigraphic unit LSF). The

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Fig. 16. (a) Upper Sarvak Fm exposed in ‘distal’ facies in NW Dome. Depositional sequences and location of logged section indicated. Upper Sarvak Fm is 270 m thick. In this location SB-Sa4 and sequence Sa5 are poorly expressed. Also note folding of Ilam Fm in hanging wall of normal fault due to ductile deformation of Surgah Fm. (b) Logged section and outcrop photograph of Upper Sarvak Fm in Central Anaran Anticline. In this region, SB-Sa4 is well developed, and overlain by an early cemented transgressive ooid-pelloid shoal complex, which is in turn capped by karst SB-Sa5. Sa5 is 40 m thick for scale. (c) Detail of glauconitic transgressive conglomerate sitting above karst SB-Sa4. (d) Karstic cavern filled with vadose silts cutting pelloidal-ooid shoal within Sa5. The caverns extend down 8 m from the karst surface SB-Sa5. (e) Detail of solution-enlarged fractures at SB-Sa5. The karst is filled by glauconitic and phosphatic silts rich in a Turonian-aged planktonic fauna of overlying sequence Sa6.

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rudists are rarely preserved in-situ, but rather occur as reworked clusters developed along foreset and cross set boundaries (Fig. 17c). (c) Platform top/interior pelloidal-miliolidChondrodonta rich micrites to wackestones with localized pelloidal packstone-grainstone tidal shoals (lithostratigraphic unit LSE, Figs 17a, b). The pelloidal-miliolid-Chondrodonta-rich facies are locally associated with laminated cherty micrites with well developed

fenestral and tepee fabrics and high-spiral gastropods. These facies are interpreted as inter to supra-tidal deposits. Spectral gamma logs (subsurface and outcrop data) indicate a significant uranium content throughout this unit, possibly reflecting proximity to a clastic source. The pelloidal packstone-grainstone shoals have well developed tidal bedforms, and are interpreted as migrating sub-tidal platform top shoals developed in slightly deeper

Fig. 17. (a– b) Outcrop photographs of Thalassinoides sp., bioturbated platform top miliolid, Dasycladacae, Chondrodonta-rich mudstones and wackestones of lithostratigraphic unit LSE. Cycle tops are characterized by erosive-based wave and current rippled packstones, often associated with a lag of disarticulated Chondrodonta shells. Pervasive Thalassinoides burrows extend off cycle tops and represent omission colonisation surfaces. The burrow fills and storm lags are preferentially dolomitized by D1 dolomite. In thin-section, D1 dolomites in the burrow fill are characterized by zoned dolomite rhombs with inclusion-rich cores and clear rims. See also Figure 31b. (c) Outcrop photograph of Radiolitid rudist floatstone-packstone facies typical of the Sarvak Fm platform margin. Finger for scale. (d) Outcrop photograph of highly bioturbated cherty slope wackestones and packstones cut by stylolites. D1 dolomites are concentrated around early cemented nodules (characterized by uncompacted burrows) and along stylolites. See also Figure 31a.

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parts of the platform top. Similar facies developed immediately above the platform margin facies (ii) are interpreted as wash-over fans and tidal deltas. In favourably orientated outcrops and seismic sections the three facies associations define a well developed clinoform geometry associated with platform/shelf margin progradation. The steeper clinoforms (3–108) are developed within the platform margin shoals (facies ii, LSF). Our outcrop observations in Lurestan indicate that Sarvak sequence Sa1 displays both accretionary low-angle ramp (,18) and steep-dipping (3 –108) clinoform shelf margin geometries, presumably reflecting basin geomorphology and windward-leeward relationships (Embry et al. 2008). Regional mapping also indicates that Sarvak sequence Sa1 has a complex progradation orientation, with both a north and south facing margin evident working away from the outcrops in the Anaran Anticline. The north facing margin is located between Kuh-E-Anjir and Kabir Kuh, whilst the south facing margin appears to have prograded southwards into the remnant Kazhdumi intrashelf basin. Progradation towards the NW, N and NE from the SE (Dehluran) and SW (Iraq) is also evident. By end Sarvak sequence Sa1, the remnant Kazhdumi intrashelf basin had clearly been filled in the Anaran area. Lower Sarvak sequence 2 (Sa2) sits transgressively over Lower Sarvak sequence 1, and comprises a very thin/condensed TST/MFS interval (2–3 m oligosteginid mudstone, MFS-Sa2) overlain initially by HST platform margin shoals followed by a thick aggradational package of platform top facies (90 m, Figs 11 & 15). The sequence is entirely of Early Cenomanian age, and is characterized by abundant benthic forams for example, Orbitolina sp. and Praealveolina sp. and limited rudistic facies. Two exposure surfaces and hard-grounds associated with reddish silt are evident towards the top of the HST, the second of which is associated with a fracture-controlled micro-karst (SB-Sa2, Fig. 10), which is abruptly overlain by an interval of high energy stacked sub-tidal packstone-grainstone shoals and subordinate platform top wackestones (unit LSA, TST of overlaying Sa3, Figs 11 & 15). SB-Sa2 is locally associated with an incised channel geometry, with overlying LST–TST charophytebearing packstones to grainstones interpreted as transgressive valley-fill deposits (cf. Top Natih E, Gre´laud et al. 2006, 2010). The lithostratigraphic top of the Lower Sarvak Fm is marked by an abrupt contact to outer platform-basinal ammonite-bearing micritic limestones related to latest Early Cenomanian MFS-Sa3. This unit is lithostratigraphically assigned to the Ahmadi Member of the Upper Sarvak Formation.

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Upper Sarvak Formation The Upper Sarvak Formation is 270 m thick at outcrop in the Anaran Anticline and 345 m thick in the subsurface to the south. It may be divided into four depositional sequences (Figs 10, 14 & 16); (a) Sarvak sequence 3 (Sa3) is of latest Early Cenomanian to earliest Late Cenomanian age. The TST at the base of Sa3 comprises transgressive valley fill or sheet-like tidal shoals (unit LSA, 30 m) abruptly overlain by ammonite-bearing outer platform-basinal micritic limestones and shales of the Ahmadi Mbr (25 m, Figs 14 & 16a). The contact between LSA and the Ahmadi Member is typically marked by a bored iron-glauconite hardground. Sa3 HST is 65 m thick at outcrop and up to 125 m thick in the subsurface to the south. It is characterized by pelagic to hemipelagic oligosteginid and planktonic foram rich wackestones and mudstones, including slope to basin floor channel-leve´e and lobe facies in the NW Dome section (Fig. 14). These facies pass into dolomitic bioclastic wackestones interpreted to have been deposited in outer to mid shelf settings in Central and Southern Anaran, which in turn pass into shelf margin facies in Siah Kuh and the Dehluran field. The HST is capped by a subaerial erosion surface (SB-Sa3, karst in platformal areas, e.g. Siah Kuh outcrops and Dehluran field; Taghavi et al. 2006, 2007) associated with significant porosity creation and secondary dolomite development. At outcrop in Anaran, SB-Sa3 is poorly expressed in slope facies, although it is typically associated with a zone of enhanced meteoric diagenesis and secondary dolomitization. Within a regional context, Sarvak Sequence Sa3 appears to be equivalent to the Rumalia Fm in Iraq, and progrades from the south towards the Garau Basin in the north. (b) Sarvak sequence 4 (Sa4) is of Late Cenomanian age (Mishrif Fm equivalent). At outcrop, it comprises a thin LST and TST associated with amalgamated turbidite channels overlain by a strongly progradational HST. Where fully developed, the HST records a thick succession from outer shelf wackestones through cherty-dolomitic bioclastic mid shelf facies up to rudist-rich margin and muddy platform top facies and a capping conglomerate and karst (SB-Sa4). In favourably orientated sections, seismic scale clinoforms (,18 to 58 dips) are evident within the shelf to rudist margin facies. In outcrops NE and E of Anaran (Rit, Chenareh and Khorram-Abad Anticlines), a detached ‘down-stepped’

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platform has been identified lateral to karst Sa4, and is interpreted as a forced regressive unit. This unit is poorly developed in northern Anaran (Figs 14 & 16). In Central-Southern Anaran a truncated sequence Sa4 is cut by karst Sa4, and directly overlain by the TST of Sarvak sequence Sa5 (Figs 10 & 16). (c) Sarvak sequence 5 (Sa5) is characterized by a basal oligosteginid-rich glauconitic mudstone and conglomerate overlain by a pelloidaloolitic shoal complex in Central and Southern Anaran (Fig. 16b & c). The shoals are transgressive (TST). No HST is preserved within Sa5, and the TST shoals are capped by a regionally developed penetrative karst which marks the Cenomanian-Turonian boundary throughout Lurestan (SB-Sa5). In the NW Dome section Sa5 is poorly expressed, with sequence Sa6 apparently sitting directly on sequence Sa4, or on a much reduced sequence Sa5 (Figs 14 & 16). Throughout Anaran karst deposits associated with SB-Sa5 fill both NE–SW and NW –SE solution enlarged conjugate joints and normal faults (Fig. 16e), indicating that the platform underwent normal faulting and flexural uplift followed by subareal erosion and karstification. This regional scale doming/flexing is also associated with an angular unconformity, partial removal of sequence Sa5, and onlap by the overlying Turonian sequence Sa6. A clastic glauconitic fill has locally been observed in the karst, including well developed dolines with depths of upto 10 m (e.g. Central Anaran and Chenareh anticlines). SB-Sa5 (which occurs in the middle of a depositional sequence), and the forced regressive interval locally developed at the top of Sa4 are attributed to the onset of deformation and flexural uplift along the NE margin of the Arabian plate. This deformation occurred from Turkey in the NW to Oman in the SE due to intra oceanic thrusting and the commencement of ophiolitic nappe emplacement, and accounts for the complex stratal architecture which typifies the latest Cenomanian-earliest Turonian throughout this region (Robertson & Searle 1990; Stoneley 1990; Burchette 1993; Homke et al. 2009). (d) Sarvak Sequence 6 (Sa6) is of Middle Turonian to Late Turonian–Early Coniacian age. Karst Sa5 is overlain by a transgressive glauconitic conglomerate and pelloidal shoals which pass rapidly up into a condensed nodular bedded interval rich in ammonites (Ghirab Member, MFS-Sa6). Dating of ammonites and microfauna from the MFS give a latest Middle Turonian age. The Early Turonian has not been identified. The MFS is interpreted as a

significant baffle/barrier horizon, and is overlain by a progradational HST characterized by mid shelf facies at outcrop and shelf interior/platform top facies in the subsurface, indicating a progradational direction towards the north. In proximal locations, the HST is capped by a karst (SB-Sa6), and abruptly overlain by claystones of the Surgah Fm (e.g. Dehluran field, Taghavi et al. 2006, and Siah Kuh outcrops).

Surgah Formation At outcrop the Turonian–Coniacian aged Surgah Fm comprises a 40 m thick claystone interval with thin micritic limestone interbeds. This interval thins to 2 m in the subsurface (SW) and thickens to over 100 m at outcrop in the NE (Kabir Kuh). The top Sarvak–Surgah formation boundary is conformable where studied at outcrop in the Anaran Anticline, and coincides with the Late Turonian– Early Coniacian boundary (Fig. 14). The Surgah Fm is interpreted as a regressive clastic wedge (LST) related to deformation of the north-eastern margin of the Arabian Plate (Setudehnia 1978; Alavi 2004; Homke et al. 2009). Chlorite-smectite rich claystones of the Surgah Formation form an excellent regional top seal to the underlying Sarvak Formation reservoirs. The Surgah Formation is also interpreted to form an excellent top seal to the fracture-fed dolomitic bodies, and an excellent clay smear unit when involved in faulting (Figs 4 & 16a).

Ilam Formation The contact between the Surgah and Ilam formations is conformable and associated with a return to carbonate deposition. An MFS (MFS-Il1) is picked at the top of the Surgah Fm/base of the Ilam Fm. The Ilam Formation can be divided into two main sequences (Fig. 10). (a) Ilam sequence 1 (Il1) comprises a thin LST/ TST, predominantly within Surgah Fm claystones and pelagic micrites, overlain by a thick HST of shallow shelf to mid-inner ramp pelloidal wacke to packstones rich in Rotalia sp., green algae, Inoceramus sp., and rudist debris capped by a karst (SB-IL1). Ilam Sequence 1 is predominantly of Middle Coniacian age. (b) Ilam sequence 2 (Il2) is of latest Middle Coniacian, Santonian to base Early Campanian age. Karst SB-Il1 is abruptly overlain by a complex of stacked and channelized wackestones and packstones interpreted as LST outer shelf to toe of slope turbidites which deepen up to MFS Il2. The overlying HST comprises a relatively monotonous succession of oligosteginid-rich marls and wackestones. The Early

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Campanian MFS (MFS-Il3) marks the contact to the overlying Gurpi Fm, and is associated with glauconite and phosphatic nodules.

Dolomites – field observations Multiple dolomite fabrics can be recognized in the studied outcrops. Our data also allows the identification of multiple dolomite stages. Fabrics and stages of dolomitization identified in the field are confirmed by detailed thin section petrographic observations, although it is important to stress that the key cement stratigraphic-structural-facies relationships were best pieced together based on observations in the field, and in particular by undertaking vertical and lateral ‘diagenetic transects’ away from fault zones. This is in keeping with the view of Warren (2000) that interpretation of dolomite genesis is often best deduced from outcrop/ external geometry. Successive and distinctive dolomite and calcite cement stages are termed D1, D2, D3, D4, D5, C1, C2. . . . etc. and are described below. Simplistically, three main stages/types of dolomitization are evident: (a) D1, fabric/facies selective. Observed relationship to stylolites. Volumetrically moderate. (b) D2, fracture-controlled, fabric selective. Volumetrically large. (c) D3-D5, fracture-controlled saddle dolomite, typically cement in D2 dolomites and prevalent along faults. Volumetrically small. D2 and D3–D5 are interpreted to be closely related in time, and are best interpreted as having developed ‘distal’ (D2) and ‘proximal’ (D3–D5) to fault zones (cf. Yao & Demicco 1995; Swennen et al. 2003). D1 dolomites could also have developed during the same dolomitization event as D2 and D3– D5, occurring ‘distal’ to D2 dolomite, although our data does not conclusively demonstrate this as yet (research in progress). D3–D5 dolomites are volumetrically more significant towards the centre of dolomite bodies and at deeper stratigraphic levels in the Anaran outcrops.

D1 matrix replacement dolomite D1 dolomites are developed within three distinct facies (Fig. 10): (i) Chondrodonta, miliolid and Dasycladacae algae rich platform interior/lagoonal micrites, wackestones and packstones. The dolomites are focused within Thalassinoides burrows and thin wave-rippled intervals (Fig. 17a, b). The Thalassinoides burrows extend off scoured bed boundaries, and typically are capped by Chondrodonta sp. storm lags

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associated with current and wave ripples. These boundaries are interpreted as storm truncated omission colonization surfaces. Dolomitization appears to have affected the slightly ‘grainier’ fill of the burrows and the overlying wave-ripple storm beds. The facies and fabrics are comparable to early dolomitized burrow fills described by Cantrell et al. (2004) from the Arab Fm in the Ghawar field and by Horbury & Qinq (2004) from the Carboniferous of Northern England. However, the fabrics are also comparable to late burial-related dolomitized burrows from the Natih Formation in the Fahud field of Oman (Vahrenkamp et al. 2006; Taberner et al. 2007). Type i facies are most prevalent within lithostratigraphic unit LSE of Lower Sarvak sequence Sa1 in the Anaran Anticline, but are also developed within Sa2, Sa3, Sa4, Sa6 and Il1 elsewhere in Lurestan (Fig. 10). (ii) Mid to upper shelf bioclastic Orbitolina-rich wackestones to packstones. Bed boundaries are often associated with wave scoured omission colonization surfaces from which robust Thalassinoides burrows extend. The burrows are dolomitized. Facies ii is present in the Mauddud Mbr and Sarvak sequence Sa1, Sa2, Sa3 and Sa4 (Fig. 10). (iii) Outer to mid shelf cherty bioclastic wackestones rich in planktonic fauna including oligosteginids, globigerinids, in-faunal echinoderms and siliceous sponge spicules (Fig. 17d). Bioturbation is intense (Bioturbation index 4/5: Thalassinoides, Planolites, Palaeophycus, Anconichnus, Chondrites, Helminthopsis, echinoderm burrows). The Thalassinoides burrows are pervasively dolomitized. This facies is very common in the Sarvak Fm, particularly sequences Sa3, Sa4 and Sa6. In all three facies, bedding-parallel to slightly oblique stylolites and stylo-cumulates are developed around early cemented nodules (characterized by uncompacted burrows). There is a clear relationship between the occurrence of stylolites and dolomite, with dolomite rhombs concentrated along and immediately adjacent to stylolites (Fig. 17). This relationship is interpreted to indicate that dolomitizing fluid flow was focused along stylolites (cf. Graham et al. 2003; Graham Wall et al. 2006), indicating that dolomitization occurred under at least some burial depth (1000 m þ ?) that is, following compaction and cementation. Petrographically, D1 dolomites can be difficult to distinguish from D2 dolomites, both characterized by inclusion-rich dolomite rhombs. Field mapping of D1 dolomites indicates that they are laterally persistent over 10s of kilometres,

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Fig. 18. (a) Outcrop photograph of D2 dolomite body, core of Anaran Anticline, NW Dome (See Fig. 5a for location). D2 dolomite is picked out by dark brown colour. Base of section (base of gorge) is 100% dolomite. Top of section is 100% limestone. Notice how vertical extent of D2 dolomite is inhibited by major bedding breaks (MFS mudstones), which are interpreted to have behaved as aquitards. In the centre of the photograph, the first major aquitard (MFS-Sa3, Top Lower Sarvak) is breached by a number of closely spaced small displacement faults and associated fractures. In this region, the D2 dolomites have extended up to the next major aquitard (MFS-Sa6), which again hindered the vertical advance of dolomite. Where breached by fractures, dolomite can be mapped to extend up to the top of the Sarvak Fm, before being stopped at the contact to ductile claystones of the overlying Surgah Fm, which act as the main top seal to both hydrocarbon migration and dolomitization. Upper Sarvak is c. 270 m thick for scale. Boxed area is enlarged in b. (b). Detail of a dolomite plume developed at the contact between the Lower and Upper Sarvak Fm (boxed area in a). At this location, a normal fault tips out upwards (looses displacement) at the contact with mudstones of the Ahmadi Member. D2 dolomite can be mapped to have developed along this fault zone, with a plume developed immediately above the fault, and multi-directional stratabound fingers developed laterally away from the fault in unit LSA. The stratabound fingers occur below cycle tops/bed boundaries. Note also bulbous irregular bases and sharp tops of stratabound D2 dolomites. The recessive-weathering non-dolomitized facies below the plume are muddy lagoonal micrites of lithostratigraphic unit LSB (sparsely jointed, low matrix poro-perm). Unit LSA is 45 m thick for scale. In both Figures a and b note doming above the D2 dolomite bodies.

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although volumetrically small to moderate in distribution in the country rock (10 –20% of total rock volume prior to D2 dolomitization). Field mapping has also clearly demonstrated that D1 dolomites were utilized as flow pathways and nucleation sites for D2 dolomites, as the spatially most extensive stratabound D2 dolomites can be ‘walked out’ in to D1 dolomites (typically Thalassinoides-burrowed intervals – discussed below).

D2 matrix replacement dolomite D2 dolomite is volumetrically the most significant dolomite in the study area and occurs as massive dark reddish brown bodies in the outcrops. The colour contrast to the light grey and white limestone country rock is striking, and it is this colour contrast that allows spectacular visualisation of the D2 dolomite geometries and relationship to fractures and facies (Figs 4, 5 & 18). D2 dolomites occur as massive, domal and plume-like bodies,

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sub-horizontal stratabound layers, and as halos to both faults and joints (Figs 18 –22). Both gradational (diagenetic front) and fracture ( joint, fault) terminations to D2 dolomite bodies are observed (Fig. 19). Vertically, D2 dolomite bodies either have gradational convex-up diagenetic tops to limestone country rock (Fig. 20b) or stop very abruptly at low permeability stratigraphic layers such as mudstones, claystones and muddy wackestones in which joints are sparse to absent (Fig. 18b). D2 dolomite bodies can be mapped up to the top of the Sarvak Formation, and are only very rarely observed above the Surgah Formation. The Surgah Formation is c. 40 m thick at outcrop, thinning to only a few metres thick in the subsurface. It comprises an argillaceous claystone rich in chlorite and smectite. Most joints and major (normal) faults developed within the Sarvak Formation stop at the base of the Surgah Formation. Normal faults loose displacement very rapidly in the Surgah claystones as displacement is taken up by ductile folding

Fig. 19. (a) Diagenetic termination of a D2 stratabound dolomite tongue. Note visible vuggy porosity development in centre of dolomite body, and lack of porosity in adjacent limestone country rock. Unit LSG, NW Dome. (b) Diagenetic edge of the Kuh-E-Pashmi dolomite body (See Fig. 20b). Bedding within turbidites of Upper Sarvak sequence Sa6 can be ‘walked’ into the dolomite body, but bedding definition becomes difficult to define in the dolomite body. Iron sulphides are developed along the edge of the dolomite-limestone contact. This diagenetic contact can be mapped out on the top and two sides of the Kuh-E-Pashmi dolomite plume. (c) Fractured (joint) edge to D2 dolomite body (white dashed line) developed in ramp interior muddy carbonates of the Gadvan Fm exposed at base of gorge in NW Dome. Joints strike NE– SW. Dolomitization is also controlled by bedding breaks, which are represented by thin bioturbated siltstones overlying firm/hardgrounds at the top of individual ramp interior carbonate cycles.

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Fig. 20. (a) D2 dolomite plume developed in footwall of a normal fault. Fault plane strikes parallel to cliff face, and trends NW (right)-SE (left) – see Figure 5 for location. Note subtle doming above plume associated with development of radial fractures. Also note stratabound D2 body extending laterally within karstic horizon LSC and below lagoonal micrites (LSB) which define top of Lower Sarvak Sequence Sa2. (b) Outcrop photograph of Kuh-E-Pashmi dolomite plume. The upper and lateral edges of this plume are diagenetic (Fig. 19b), whereas the lower lateral edges are in part fault bounded. Note massive nature of core of dolomite body, and absence of bedding, which can be traced into well bedded and jointed limestones of Upper Sarvak sequence Sa6 (right edge of photo). Also at this locality organic-rich mudstones associated with MFS-Sa6 appear baked at the contact to the dolomite body. A transect at the MFS level (white dashed line) reveals bitumen and local hydrocarbon adjacent to the dolomite, particularly developed along joints. (c) Stratabound dolomite developed within lithostratigraphic unit LSG of the Lower Sarvak sequence Sa1. Note sharp top (defined by bioturbated bed boundary) and irregular bulbous lower boundary (white arrows). Massive dolomite in lower left part of photograph is developed along top Mauddud karst (SB-Kz3), but extends vertically in a fracture corridor (also visible on D). (d) Example of stratabound dolomites developed in top Mauddud and base Lower Sarvak. Dolomite follows Sb-Kz3 and D1 dolomitized burrow horizons. More extensive vertical dolomite distribution in right side of image is related to a fracture corridor.

(Fig. 16a). Where faults do cut through to the overlying Ilam Formation the Surgah Formation typically forms a very well developed smear along the fault zone. These observations indicate that the Surgah Formation would have been a very effective top seal to any vertical movement of dolomitizing fluids, which is also in keeping with the role of the Surgah Formation as the main top seal to hydrocarbon reservoirs of the Sarvak Formation in the nearby subsurface. On the scale of the studied river gorge in the NW Dome area of Anaran Anticline (c. 8 km long by 1 km deep section), D2 dolomites are volumetrically more abundant in the deepest exposures, and decrease in volume successively upwards (Figs 5 & 18). Comparable observations can be made in

several other sections in the Anaran and adjacent anticlines (e.g. Kuh-E-Pashmi, Anjir Anticline, Rit Anticline). Significant upward volume reductions in D2 dolomites are evident at major stratal contacts as listed below: † Base of the Surgah Fm (SB-Sa6) † Base of the Ghirab Member (MFS Sa6) † Thin argillaceous transgressive mudstone capping SB Sa4 † MFS Sa4 † Base of the Ahmadi Member (MFS Sa3 zone) † Thin argillaceous transgressive mudstone capping SB Sa2 † MFS Sa2 † Thin argilaceous mudstone above Mauddud Mbr karst (SB-Kz3)

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Fig. 21. D2 dolomites and fault zones. (a) D2 dolomite developed along damage zone/fault breccia in hanging wall of a normal fault extending above the Kuh-E-Pashmi dolomite body. Offset is c. 15 m. The fault trends parallel to the elongation of the Anaran Anticline (SE–NW, F1 faults). Several such dolomitized faults can be mapped in the Kuh-E-Pashmi area extending from the main dolomite body. (b) Dolomitized NE– SW trending (F2) normal fault in the Kuh-E-Pashmi area. The outer (most recent) slip surface of this fault is associated with re-brecciation of the D2 dolomitized fault breccia, with the younger breccia comprising D2 dolomite clasts cemented by calcite (inset photo). This indicates a change in mineral chemistry of the fluids circulating along the faults during the late stages of deformation. Map board (30 cm long) for scale. (c) Normal fault zone cutting Lower Sarvak Fm in NW Dome. Fault plane and breccia are partially dolomitized. Boxed are is enlarged in D. (d) Detail of fault breccia developed along fault shown in C. Host rock clasts remain as unaltered limestone, whereas the permeable fault rock is preferentially dolomitized. Hammer head for scale.

† MFS Kz2 † MFS Kz1 † MFS Da2 All of these boundaries are associated with distinct low matrix porosity-permeability and low fracture ( joint) density mudstone/argillaceous facies (Fig. 10). Joints and faults also often stop at these mudstone/argillaceous breaks, indicating that they represent major breaks in the mechanical stratigraphy. Where joints and faults have cut these muddy/argillaceous intervals, D2 dolomite bodies can be mapped to extend upwards along the fractures and extend laterally in more permeable/ jointed facies of the overlying sequence, before being stopped or retarded by the next argillaceous layer. This situation is spectacularly exposed in

Figure 18. These data clearly indicate that the low matrix porosity and low fracture ( joint) density units hindered or stopped the vertical migration of dolomitizing fluids. Where D2 dolomites have been mapped along faults which tip out (i.e. loose displacement) at major argillaceous breaks, such as the base of the Ahmadi Member, subtle low-relief dome-like geometries are evident (Figs 18b & 20a). At least two of these domes have been mapped associated with fractures which radiate away from the dome centre. Sub-horizontal stratabound sheets of D2 dolomite are concentrated below the major argillaceous boundaries listed above, following specific facies or bed boundaries in the underlying units (grainstones, thin mudstone breaks or cemented cycle tops). The tops of these stratabound sheets typically

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Fig. 22. Bedding plane outcrop of jointed Turonian mudstones associated with MFS-Sa6. Both NE–SW and NW–SE trending joints are well developed, as well as en-echelon shear fabrics. NE is towards page top. The joints are associated with microporosity development and partial dolomitization, picked out by light brown coloured ‘halos’ along the joints. Notice that alteration is most extensive where joints intersect. The shear fabrics have a central fill of blocky calcite. Jacobs staff is 1.5 m long, and orientated NW– SE.

are very sharp and follow lithological boundaries, whilst the bases often are very irregular and bulbous (Fig. 20c). Vug development (fabric and non-fabric selective) is prevalent immediately below cycle-capping permeability barriers. When traced laterally, the strata-bound sheets have two types of termination: (a) Diagenetic front. Lateral termination of the D2 stratabound dolomites are relatively abrupt, with 100% dolomite passing into 100% limestone in a matter of a few cm (Fig. 19a, b). Occasionally, scattered dolomite rhombs of D1 and D2 affinities are observed in the limestone country rock. Strikingly, the D2 dolomites can be associated with well developed mouldic, intercrystalline and vuggy porosity, whilst the immediately adjacent limestone country rock appears tightly cemented (Fig. 19a). (b) Fracture. Fracture controlled edges to D2 stratabound dolomites are also observed, although less common. In Figure 19c, 100% dolomitized units of the Gadvan Formation with well developed intercrystalline, mouldic and vuggy porosity can be traced laterally away from a major fault zone to where they abruptly stop at vertical NE–SW striking joints. Across the joints the country rock is undolomitized and lacks porosity development. The implication

is that the joint acted as a barrier to the lateral advance of dolomitizing fluids. The lateral extent of D2 stratabound dolomites away from the interpreted input points (faults) is strongly influenced by and follows the pre-dolomitization stratigraphic and diagenetic framework. The list below records in approximate order of importance intervals most susceptible to D2 dolomitization (see also Fig. 10): (a) Karstified horizons (e.g. SB-Da1, SB-Da2, SB-Kz3, SB-Sa2, SB-Sa5). (b) Intervals immediately beneath cycle capping mudstones, from the metre scale within individual platform top cycles, to the sequence scale beneath major transgressive surfaces and maximum flooding surfaces. (c) The tops of highstand deposits, for example Sarvak sequences Sa1, Sa2, Sa3 and Sa4. Petrographic studies indicate that cycle-capping exposure surfaces at the top of these sequences are associated with early meteoric porosity creation. These leached intervals were later utilized as flow pathways by D2 dolomitizing fluids. (d) Beneath seismic-scale mid shelf clinoform boundaries, particularly in Sarvak Fm sequences Sa1, Sa3 and Sa4 in the Lurestan area. This appears to reflect early cementation

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of set boundaries, and focus of fluid flow beneath them. (e) Porous/permeable ‘grainy’ facies, for example rudist-filled tidal channels and shoals developed within the TST of Dariyan sequence Da2. (f) Bioclast-rich facies, associated with the creation of biomouldic porosity due to dissolution of calcitic/aragonitic bioclasts. Interestingly, this does not always appear to be applicable to rudist-rich facies (e.g. within Sa1 and Sa4 in Lurestan). The reason for this is unclear as yet (research in progress). In contrast, D2 dolomite advances only minor distances away from faults in the following facies: (a) Low porosity/permeability and low joint density argillaceous mudstones, either capping small scale cycles or major MFS intervals capping sequences. (b) Low porosity/permeability slope or platform top micrites and wackestones. (c) ‘Grainy’ pelloidal/oolitic non-skeletal TST units (e.g. lithostratigraphic units KZB, LSA, USB). These units are preferential cemented by early marine cements and only moderately jointed by widely spaced joints. This combination of low matrix porosity and widely spaced joints is interpreted to have limited the access of D2 dolomitizing fluids to significant rock volumes. When all these factors are combined, a preference for D2 development at porous cycle tops (cycle, parasequence, HST sequence, karst) immediately beneath aquitards (mudstone breaks, flooding surfaces, transgressive surfaces, maximum flooding surfaces) is clear. This template is similar to that proposed by Swennen et al. (2003) and Davies & Smith (2006) for High Temperature Dolomites (HTD) in North America. Importantly, in the Iranian case study, these conclusions have allowed the establishment of a powerful stratigraphic/ diagenetic framework that may be used as a predictive tool to establish likely D2 dolomite distribution in the nearby subsurface and populate reservoir models. However, as mentioned above and discussed later, it is perhaps even more important to superimpose a fracture (joint) stratigraphy on the stratigraphic framework to fully ascertain a unit’s susceptibility to dolomitization. The importance of karst as a flow pathway during D2 dolomitization is best demonstrated at the karst horizon which caps the Mauddud Mbr (SB-Kz3, Figs 13 & 20c, d). The karst is associated with significant permeability development (solution breccias, cave layers, solution enlarged faults and joints), and is capped by a thin (5–20 cm thick) mudstone layer. D2 dolomitizing fluids utilized the

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high permeability karst, resulting in a stratabound dolomite body which can be mapped laterally over 100 km. D2 dolomites are thickest within the karstified interval, and have irregular and bulbous bases. The thickness of the interval that is dolomitized beneath the karst thins when mapped away from major fault zones, implying that the dolomitizing fluids were sourced from the fault zones. Above the karst dolomitization is limited to a thin transgressive interval of rudist shoals and grainstones, and generally does not extend up into the overlying low permeability wackestones of the Lower Sarvak Formation (unit LSG, Figs 13 & 20c, d). Exceptions to this are where major joints and faults cut through the stratigraphy, and dolomite can be mapped upwards along the faults and then outwards along the next permeable layer or beneath the next major aquitard (Fig. 18a). On a smaller scale, intensively bioturbated and dolomitized (D1) intervals (Fig. 17) are also utilized as flow pathways for D2 dolomites, with stratabound ‘patchy sheets’ of D2 dolomite extending the most significant distance laterally away from faults along these horizons. Good examples of this can be walked out within platform top facies of Lower Sarvak sequence Sa1 and mid-upper shelf facies of Upper Sarvak sequence Sa4. At the deepest stratigraphic levels in the NW Dome outcrops (e.g. Dariyan Fm) D2 dolomitization is total, although primary facies may still be identified. The originally more ‘grainy’ facies (shoal fore- and top-sets, rudistic beds, Thalassinoides intervals) form the coarser crystalline dolomites. These are also the intervals associated with the higher visible porosity at outcrop. Rudist-filled channels at the base of sequence Da2, for example, appear to have been utilized as a flow horizon. In contrast, the originally muddier shoal toesets and lagoonal micrites to wackestones form micro-crystalline dolomites with minimal porosity enhancement. The process by which D2 dolomitization of large areas of low permeability wackestones and mudstones occurred was initially unclear, for example lithostratigraphic units KZC, KZA, LSG, USD and NNS (Fig. 10), which have been observed totally dolomitized in the Anaran Anticline. Detailed observations at several well exposed outcrops outline the importance of jointing in this process. Figure 22 shows a bedding plane of jointed interbedded calcareous mudstones and pelagic micrites of the Middle Turonian Ghirab Member exposed at Kuh-E-Pashmi. The Ghirab Member represents a major MFS and a significant aquitard to the advance of D2 dolomites. In the Kuh-E-Pashmi area, however, this interval is dolomitized, with a spectacularly exposed dolomite plume developed which can be mapped up to the contact with the overlying Surgah Formation (Fig. 20b). The upper

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and lateral edges of this plume can be ‘walked-out’, and the contact is a well defined diagenetic front (Fig. 19b). The front is associated with a thin zone of iron sulphides. Moving into the immediately surrounding undolomitized country rock, halos of dolomite and microporous limestone can be mapped along subvertical joints (Fig. 22). The joints trend NE –SW and NW –SE, and there is evidence for dilational shear on both joint sets. Both sets of joints are associated with dolomite and microporous limestone, but alteration is generally more intense along the NE –SW trending set. We interpret these observations as indicating that the NE–SW trending set were optimally oriented for diagenetic fluids (i.e. oriented parallel to the maximum horizontal stress and thus open). Close examination of the joint edges allows mapping of successive diagenetic fronts which parallel the joints. In thin-section, a front of dolomite rhombs and microporous limestone is evident defining the contact to unaltered country rock. Where NE–SW and NW–SE trending joint sets intersect, the area affected is volumetrically larger, and the diagenetic bodies converge to produce a larger irregular body. Where joints are closely spaced, the total area affected is large, but where joints are widely spaced only limited areas of the country rock are affected. These small scale observations clearly indicate the importance of joints, joint density and joint aperture/permeability within specific lithologies as a primary control on flow of diagenetic fluids. If up scaled to a regional scale, we feel this process may adequately explain how large areas of low permeability rock are dolomitized. D2 dolomite bodies are intimately related to faults (Figs 18b & 21). The greatest vertical extent of D2 dolomites is always along fault zones or intensely jointed fracture corridors (incipient faults?). In both the Kuh-E-Pashmi and NW Dome areas steep dipping and resistant weathering sheets of D2 dolomite can be mapped radiating away from the tops of dolomite plumes following faults or fracture corridors (Fig. 21). Both NE–SW and NW –SE trending faults are associated with D2 dolomites. The faults are associated with well developed breccias. Figures 21c, d shows an example of a fault breccia where the individual clasts of limestone sit in a D2 dolomitized fault breccia. When mapped downwards, both the clasts and breccia are dolomitized, and when mapped upwards first the clasts are undolomitized, and then the breccia matrix. These relationships are interpreted to indicate that D2 dolomitizing fluids were flowing through the permeable fault breccia. On other faults, notably NE– SW trending faults, resistant weathering dolomitized ribs and sheets of fault rock are themselves brecciated and occur as clasts set within a calcitic fault rock

matrix (Fig. 21b). D2 dolomite bodies are also clearly offset by both NE –SW and NW–SE trending faults (Figs 4 & 5). The implication is that the flow of dolomitizing fluids along faults progressively ceased to be active through time, with D2 dolomite bodies pre-dating the last phase of faulting. In NW Dome and elsewhere in the Anaran Anticline, NW– SE trending conjugate faults which brecciate and offset D2 dolomite bodies record a rotated geometry, with NE-dipping faults rotated to shallower angles and SW-dipping faults rotated to steeper angles (Fig. 4). This indicates passive folding of the conjugate faults during the later stages of fold growth (cf. Sharp et al. 2000). Although the timing and evolution of jointing, faulting and folding are interpreted to be closely related, the detailed field observations of D2 dolomites allows the following simplified chronology to be developed: (a) Pre-folding jointing of the country rock by NE –SW and NW–SE trending joint sets. Joints develop at 908 to bedding. (b) Normal faulting of the country rock associated with the development of D2 dolomite bodies. Dolomitizing fluids flowed along faults, joints and favourable horizontal flow pathways or beneath aquitards. (c) Continued faulting and folding, but with a switch from dolomite-rich fluids to calcite-rich fluids. (d) Final stages of folding resulting in passive rotation of D2 dolomite bodies and NW –SE orientated conjugate normal faults. NE –SW trending faults continued to be active at this time, and offset NW– SE trending faults.

D3 – D5 fracture and vug cementing saddle dolomites Cement phases D3–D5 comprise very distinctive coarse crystalline white saddle dolomite cements at outcrop. Petrographically, only three saddle dolomite events are easily distinguished (D3, D4, D5), but in the field locally up to nine saddle dolomite cement phases can be identified, typically in the immediate hanging wall of normal faults (Fig. 23). The saddle dolomites are volumetrically not as significant as the D2 dolomites, and are spatially closely related to fault zones and to the deeper stratigraphic/structural levels in the outcrops. The saddle dolomites either cement dilatational ‘floating clast’ breccias adjacent to faults (crackle and mosaic fault breccia fabrics of Woodcock & Mort 2008), or fill inclined or bedding parallel shear zebra fabrics (Swennen et al. 2003; Vandeginste et al. 2005; Davies & Smith 2006), which can be mapped back to fault zones. Saddle dolomite cements lining

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Fig. 23. (a) Outcrop photograph of normal fault zone exposed in NW Dome (see Fig. 5 for location). D2 dolomites in the hanging wall of the fault are associated with jig-saw chaotic, mosaic and crackle breccias, with individual D2 clasts cemented by D3– D5 saddle dolomite. Floating clast chaotic breccias can be walked out to clast support mosaic and crackle breccias and zebra fabrics by taking a transect away from the fault along the river valley (left edge of photo). Note development of vuggy and cavernous porosity (arrowed) in immediate hanging wall. In this example upto nine separate botryoidal saddle dolomite cement phases can be identified lining the cavern wall. (c) Also evident in the footwall of the fault is the contact between the Gadvan and Dariyan Fm. Organic rich shales (TOC up to 5 %) and distal deltaic sands of the Gadvan Fm are ‘baked’ in the immediate footwall, with joints in the shales impregnated with bitumen. When traced laterally away from the fault (towards right in footwall), this baking disappears. (b) Detail of exploded jig-saw (chaotic to mosaic) fault breccia, with clasts of D2 dolomite cemented by multiple saddle dolomite cements (D3–D5). Note ‘fitted’ nature of some clasts, and embayed and dissolved edges of other clasts, implying a combination of hydraulic brecciation and corrosion of the country rock. (c) Detail of nine phases of D3–D5 saddle dolomite cements in cavern arrowed in A. Two layers with iron sulphides are evident, as are hydrocarbon fluid inclusions in the D3– D5 saddle dolomites in thin-section.

vugs and filling intercrystalline porosity in D2 dolomites are also common. The field relationships clearly point to a fault link/derivation for the saddle dolomite generating fluids (cf. Swennen et al. 2003; Davies & Smith 2006). To spatially understand the distribution and fabric development of the saddle dolomite cements, dip transects working away from major normal faults (displacements between 50 –250 m) were undertaken. Well exposed transects are located within the Gadvan and Dariyan formations at the base of the NW Dome gorge section. Figures 23– 25 show outcrop photos of our type location for

saddle dolomite textures and cement stratigraphy. In this region, a NW-SE striking, NE dipping normal fault (558 fault plane dip) with upto 50 m of displacement can be observed juxtaposing the Gadvan and base Dariyan formations in the footwall against the middle part of the Lower Dariyan Formation in the hanging wall. Outcrops in the immediate hanging wall are characterized by dilational breccias (cf. crackle, mosaic and chaotic breccias of Woodcock & Mort 2008), comprising floating clasts of D2 dolomites within a cement of D3 –D5 saddle dolomites. The D2 dolomite clasts range from truly angular crackle fabrics, to clasts which

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Fig. 24. Representative fabrics associated with development of D3– D5 saddle dolomites. (a) Domal zebra shear fabric. Sense of shear is into the page. Compare to examples from Canada and Egypt (Fig. 37). (b–d) Bedding-parallel and elongate saddle dolomite filled vugs and zebra fabrics. B is looking down onto a bedding plane. C and D are bedding parallel. (e) En-echelon shear zebra fabric filled by composite saddle dolomite cements. Shear is away from the fault zone in Figure 23. (f) Angular jig-saw/mosaic clast breccia, with cement of saddle dolomite. (g) View looking vertically down onto shoal-top vugs shown in Figure 25. Cycle bounding D2 dolomitized muddy facies (left) underlain by dissolved and D3-D5 saddle dolomite cemented vugs at top of shoal. (h) Dissolutional/corroded fabric developed in D2 dolomites, with resultant porosity filled by D3 –D5 saddle dolomites. Also notice saddle dolomite filled bi-lateral shear fabrics (sensu Davies & Smith 2006). (i) Dissolutional fabric developed towards top of shoal cycle shown in Figure 25.

are more smooth edged in appearance and have corroded margins. Both clast- and matrix-(cement) support fabrics are evident. The clast-support (crackle-mosaic) fabrics are interpreted to be related to limited dilation (e.g. Fig. 24f), whilst the matrix/cement support (mosaic-chaotic) floating clast fabrics (e.g. Fig. 24g, i) are interpreted to be related to either significantly more dilation and/or dissolution of the D2 clasts. Which process is more prevalent can be judged by examining if the D2 clasts are angular, and can be fitted together (dilation), or if they have rounded and embayed edges (dissolution). Vuggy and cavernous porosity is evident where D3–D5 saddle dolomite cements have not fully cemented the porosity created (Figs 24 & 25). In one of these caverns (1 m in

diameter) up to nine successive phases of botryoidal saddle dolomite cements could be identified in addition to two zones associated with iron sulphides (Fig. 23c). Moving further away from the fault in a dip section into the hanging wall, the chaotic, mosaic and crackle breccias decrease abruptly in volume, from 20 m thick immediately adjacent to the fault to 5 m thick 15 m away. A progression from chaotic to mosaic to crackle breccias is also evident. Moving still further away, the breccia horizons pass into several thin intervals associated with saddle dolomite filled vugs concentrated below major bedding breaks. Well developed inclined zebra sheer and domal fabrics directly comparable to those described by Swennen et al.

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Fig. 25. (a & b) Tidal shoal developed in Lower Dariyan Fm, gorge base, NW Dome. B is in-situ outcrop and A is a fallen block of the same unit which exposes the D2 and D3–D5 fabrics well. Shoal is c. 8 m thick (2 m Jacob staff in A for scale). In both outcrops muddy bottomsets pass up into increasingly grainy foresets and topsets capped by a 10 cm thick mudstone horizon (white dashed line in B) which caps the depositional cycle. Bedding-parallel zebra and shear fabrics and vugs partially filled with D3 –D5 saddle dolomite are prevalent in the upper part of the cycle, especially immediately below the capping mudstone. (c) Detail of zebra fabric developed below a bedding break (boxed area in A). Zebra fabric and saddle dolomite cements are prevalent towards the bed top.

(2003), Davies (2004), Davies & Wendte (2005), Vandeginste et al. (2005), Davies & Smith (2006), Roure et al. (2005) and Smith & Davies (2006) are also developed, typically immediately below bed tops. The sense of shear is away from the fault zone. In dip sections, the zebra fabrics are inclined systematically as en-echelon shear fabrics (Fig. 24e). In fault parallel sections domal fabrics are evident, associated with bilateral shear microfractures and ladder fractures (Fig. 24a, cf. Davies 2004, his Figs 5–15, Davies & Smith 2006). In other beds the zebra fabric is sub horizontal to bedding and associated with vug development (Figs 24b –d). The horizontal zebra fabric is clearly developed by parting (vertical separation) of individual original bed laminae and precipitation of saddle dolomite cements, as the zebra fabric can be walked out into undisturbed bedding. Individual laminae can also be visually pieced back together like a jig-saw when saddle dolomite filled zebras are removed (cf. Zebra dolomites from Western Canada, G. Davies, pers comm., 2006). The horizontal zebra fabrics can be ‘walked out’

from the fault, from dilational and sheared breccias adjacent to the fault, to inclined zebra fabrics, to horizontal zebra fabrics and then to undisturbed bedding in D2 dolomites. These lateral fabric changes are thus best interpreted as an evolutionary continuum, with the dilational chaotic breccias evolving into mosaic and crackle breccias and then into inclined/sheared zebra fabrics which in turn evolve into horizontal zebra fabrics which in turn pass into isolated bedding-parallel saddle dolomite filled vugs and ultimately undisturbed bedding. Despite the intense alteration, primary depositional fabrics can be identified in the hanging wall outcrops of the Lower Dariyan Fm. Figure 25 shows a well developed platform top shoal cycle, characterized by muddy sub-horizontal nodular and bioturbated bottomsets passing up into inclined grainy foresets and relatively flat lying topsets and a thin capping mudstone (5 cm thick). The entire shoal cycle is 8 m thick and has been altered to D2 dolomite. D3–D5 dolomite is most pervasive in the upper part of the shoal, with closely spaced, partially saddle dolomite cemented vugs prevalent

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immediately below the capping mudstones. Saddle dolomite cements and shear fabrics are well developed within the foresets, and individual zebras have maximum dilation in the upper part of the sets and loose dilation and angle as they pass down into the muddy toesets. In the second detailed studied fault example (Figs 26 & 27), developed within platform top facies of Lower Sarvak sequence Sa1, an almost identical fabric progression to that developed in the Dariyan Fm described above could be identified. Moving away from a normal fault zone in the immediate hanging wall, dilational chaotic, mosaic and crackle breccias cemented by saddle dolomite can be mapped to pass into a series of inclined arcuate shear saddle dolomite veins (zebra fabric) which progressively flatten out into a horizontal

Fig. 26. (a & b). D3–D5 fabric development in the immediate hanging wall to a 250 m displacement normal fault cutting Lower Sarvak unit LSE in NW Dome. The fault occurs just to right of photo in A. In the hanging wall a series of 1– 2 m thick platform interior cycles are developed capped by cemented and stylolitic hardgrounds (finger is on contact in A). Beneath the hard-ground in A, an exploded chaotic jig-saw breccia adjacent to the fault (detail in B) passes in to inclined shear zebra fabrics moving away from the fault (A, left centre). This fabric progression is comparable to example in Figure 20, but on a smaller scale. Also note apparent geopetal fabric to right of clast labelled D2 in B.

zebra fabric which also thin and then terminate away from the fault. This lateral fabric progression is evident in excess of 10 cycles in the immediate hanging wall of the fault (Fig. 27). The fabrics are contained on either a bed by bed (1 m thick) or depositional cycle (3–5 m thick) scale (Fig. 27). The lateral extent of the fabrics away from the fault in this case is relatively limited (1–10 m), despite a fault displacement of 250 m. Partially cemented vugs with saddle dolomite are almost always located towards the tops of individual beds/cycles, with both fabric-selective and non fabric-selective vugs being evident (Figs 26 & 27). Bed tops are associated with mudstone breaks with horsetail stylo-cumulates. The breccias are again characterized by D2 dolomite clasts set in a white cement of D3–D5 saddle dolomite. Embayed and irregular edges to D2 clasts indicate a period of corrosion prior to precipitation of D3–D5 cements (confirmed by petrographic study). Geopetal crystal fills are also locally present (Fig. 26, cf. Davies & Smith 2006). The breccia and zebra fabrics are predominantly developed in the upper part of radiolitid rudist– Chondrodonta shell beds which are capped by thin red-stained mudstone layers interpreted as minor cycle top exposure surfaces (Figs 26 & 27). Both rudists and Chondrodonta shells are mimetically replaced by D3–D5 saddle dolomite, whilst the matrix comprises D2 dolomite with visible intercrystalline porosity. Microfractures filled with saddle dolomite cross-cut the matrix and link cements in both vugs and mimetically replaced shells. At this location, vugs and fractures have a central fill of calcite and red-pink silt, which are the youngest cement phases identified at outcrop (Fig. 26). Several normal faults can be ‘walked out’ up section in the NW Dome outcrops, from the gorge base to top (almost 1 km of section). A clear vertical diagenetic trend is evident in these vertical transects. Simplistically, the deepest outcrops have the most complex petrographic history, with up to nine stages of saddle dolomite (D3–D5) being evident in the field, often associated with the most complex tectonically-related textures and fabrics (dilational jig-saw, ‘floating clast’ chaotic-mosaiccrackle breccias and zebra shear fabrics). Moving up section, the number of identifiable saddle dolomite (D3– D5) cement stages decreases, associated with a corresponding reduction in the degree of brecciation and zebra fabrics. Towards the top of the section, saddle dolomites are absent or only locally identifiable, and D2 dolomites form the most pervasive dolomite. At the very top of the section, D2 dolomites also abruptly reduce in volume. Similar petrographic relationships are evident in a transect along the river gorge of NW Dome from the core of the Anaran Anticline

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Fig. 27. Log and photos of dolomite fabrics developed in platform top facies of Lower Sarvak unit LSE, NW Dome. Same location as Figure 26. Note preferential development of non-fabric selective vugs immediately beneath cycle capping exposure surfaces, hardgrounds and associated mudstones. Vugs are partially filled by D3–D5 dolomite and rare fluorite. Bedding parallel and domal zebra fabrics are also well developed locally. Lower right photograph illustrates non-fabric selective D3– D5 filled vugs linked by saddle dolomite cemented hairline fractures. Late calcites and geopetal sediment are developed in the centre of these vugs. Calcite and pink vadose silt-filled fractures are the latest petrographic fill, and cut the D2 and D3 –D5 dolomites.

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towards the SW flank. The core has the largest volume of D2 and D3–D5 dolomites, whilst the SW flank has limited volumes (Fig. 5). Solid bitumen and live oil have been observed at outcrop in the normal fault zones in NW Dome. These hydrocarbons are most prevalent in the deepest stratigraphic sections and are typically (although not exclusively) associated with the development of D3 –D5 dolomites. Fluid inclusion study confirms the association of hydrocarbon with the D3 –D5 and late calcite cements. Bitumen and oil impregnation are particularly evident where the Gadvan Fm organic-rich shales (TOC 3–5%) are exposed adjacent to normal faults (Fig. 23). The shales have a ‘baked’ (i.e. weakly metamorphosed) appearance and are very light in weight. Bitumen is developed along joints in the shales. When these organic-rich shales are traced horizontally away from the fault zone (less than 50 m laterally), bitumen impregnation decreases and the ‘baked’ appearance is progressively lost. Analysis of organic matter maturity and vitrinite data of the shales away from the fault indicate maximum temperatures of 70–808 and burial depths of 2.5 km. In contrast, organic matter maturity, fluid inclusion and vitrinite samples taken adjacent to the fault zone indicate temperatures of up to 1308. This temperature gradient occurs in the same bed over a horizontal distance of less than 50 m from the fault zone. A similar situation was encountered at Kuh-E-Pashmi, where TOC-rich mudstones associated with the Turonian MFS-Sa6 showed the appearance of bitumen and live oil when traced into the dolomite body (Fig. 20b). These data are interpreted as evidence for elevated heat flow along the fault zone and associated local maturation of organic rich layers where they are juxtaposed against the fault. Research in progress is testing this hypothesis. In summary, our outcrop observations clearly indicate that D3– D5 saddle dolomites are spatially related to fault zones, forming cements within fault breccias or shear fabrics which show a sense of shear away from individual faults (Figs 28 & 29). Secondly, the observation that saddle dolomites, zebra fabrics and dilational breccias are volumetrically most abundant in the core of the anticline at the stratigraphically deepest exposures, and show a progressive decrease both stratigraphically upwards and away from the core of the anticline towards the flanks clearly indicates that they are sourced from depth. Thirdly, the observation of local ‘forced maturation’ (cf. Davies 2004) of organic rich shales and temperature anomalies adjacent to fault zones and in association with the D3–D5 saddle dolomite cements points to elevated heat flow along the faults. That is, the presence of hydrothermal systems.

Our outcrop observations are directly comparable to the observations of Swennen et al. (2003) and Davies & Smith (2006), who proposed a model for HTD development, and in particular saddle dolomite cements, along faults during co seismic deformation in an overpressured system, resulting in hydrofracturing, brecciation and the development of zebra shear fabrics (Figs 5– 15, Davies 2004; Davies & Smith 2006).

Late calcites and vadose sediment The last cement phases observed at outcrop are vug, fracture and cave filling calcites and pink silts/muds (Figs 26 & 27). The calcites fill remaining porosity in both the D2 and D3 –D5 dolomites. Thick polyphase calcitic cements are also developed along fault zones, typically cementing fault rock which comprises D2 and D3– D5 dolomites. These observations indicate that faulting continued after development of D3–D5 saddle dolomites, but that a change in fluid composition circulating along the faults occurred. Fluid inclusions from the calcites confirm continued elevated heat flow along the faults (80–130 8C). The last cement phase comprises friable pink silts and muds, which are often developed as geopetal fills in vugs, fractures and caverns. These sediments are best interpreted as exposure-related terrestrial vadose sediment.

Cement stratigraphy Figure 30 shows a summary of petrographic, fluid inclusion and isotope data for the study area. As this paper focuses on the documentation and description of field geometrical relationships, only a summary of the salient petrographic data are included here. A full petrographic study of the Anaran dolomites by Lapponi et al. was prepared for the EAGE Arabian Plate Workshop, Jan 2010 (to be published by the Geological Society, London in due course).

Shallow to burial diagenesis Shallow burial diagenesis is characterized by aragonite dissolution (creating mouldic porosity), a suite of early marine calcite cements (EC1-EC3), fractures and local dolomitization. Early marine equant calcite cements are prevalent in grainy facies, particularly in transgressive non-skeletal grainstone intervals, for example, lithostratigraphic units LSA, USB (Fig. 10), and fill mouldic porosity after aragonite. Secondary chert, locally mimetic of bioclasts, also develops in mid slope facies, as do small calcite-filled pinnate hydrofractures at the

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Fig. 28. Schematic block diagram based on field observations summarizing D1, D2 and D3– D5 fabric development ‘proximal’ and ‘distal’ to fault zones. Sketched block represents hanging wall to a normal fault which bounds the bottom right hand edge of the block. A fabric progression from a complex polyphase D2 and D3–D5 cement stratigraphy adjacent to the fault to a less complex cement statigraphy and fabric development away from the fault is shown. Note the evolution from chaotic jig-saw breccias, to mosaic and crackle breccias, to inclined zebra fabrics, to bedding parallel zebra fabric, to isolated vugs, to D2 and D1 dolomite moving progressively away from the fault zone.

378 I. SHARP ET AL. Fig. 29. Conceptual summary figure of D1, D2, D3– D5 field relationships based on outcrops in the NW Dome and Kuh-E-Pashmi region of the Anaran Anticline. Vertical section is c. 1 km, and cross section is c. 3 km. D2, D3– D5 dolomites are spatially linked to fault zones (fracture fed vertically), and follow below aquitards or within permeable facies (karst, HST units) laterally. Top seal to the whole system is the clay-rich and ductile Surgah Formation.

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Fig. 30. Summary of the paragenesis for the Khami and Bangestan Groups in the Anaran Anticline. Depth is on left vertical axis, time on horizontal axis. Porosity increase and decrease schematically shown at base. EC – Early calcite. LC – Late calcite. D – Dolomite. F – Fracture phase. Diagenetic phase i is the early diagenesis (Early marine calcite cements EC1 þ EC2 þ aragonite dissolution) to burial phase. Burial is associated with calcite cements (EC3) increasing temperatures, compaction (stylolites), and cumulates in the onset of deformation at ca 8 Ma and D1 to D2 dolomitization associated with fracturing. Dolomite phase ii is associated with proven hydrothermal fluids (fluid inclusion temps of 80– 120 8C) and hydrocarbon generation, as is late calcite stage iii. Phase ii is also associated with several dolomite replacement and dissolution events. Phase iv is related to uplift and unroofing of the Anaran Anticline and a progressive cooling of the system.

edges of early diagenetic uncompacted nodules which are bounded by stylolites. D1 dolomites are interpreted as predominantly related to burial stage diagenesis based on textural relationships observed in the field and in thin section. Petrographically, D1 dolomites are finely crystalline, euhedral dolomite rhombs, which are usually minor in abundance but locally completely dominate samples. D1 dolomites typically occur as dolomitic limestones. Rhombs are replacive, both of matrix and wall-structures of foraminifera. Stylolite/pressure solution development occurs during burial, and D1 dolomite is preferentially developed along stylolites in muddy carbonate facies (basin/ slope or platform top/lagoon), suggesting syn- or post-stylolite formation (Figs 17 & 31a, b). The concentration of D1 dolomites along stylolites points to a strong permeability control on precipitation (cf. Graham et al. 2003; Graham Wall et al. 2006). D1 dolomites also developed within grainy Thalassinoides burrow fills, indicating that

depositional fabric is important in controlling dolomitization. D1 dolomitization can locally be observed in close association with early marine calcite cements and porosity destruction, raising the possibility that at least some D1 dolomites precipitated at shallow burial depths. Three suites of early fractures also are attributed to early burial, near surface diagenesis. These include compactional stylo-fractures associated with the formation of stylolites, early crumbly-edged fractures related to bioturbation and sediment movement, and en-echelon fractures attributed to tectonic fracturing during the early stages of burial. These fractures are almost always cemented by early calcite cements, and pre-date D1 dolomites.

Late burial diagenesis D2 dolomites range from microcrystalline dolomites to coarse crystalline dolomites, typically with inclusion rich centres and clear outer rims.

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Fig. 31. Thin-section photomicrographs representative of D1, D2, D3– D5, late calcite and vadose silt cements. All sections are impregnated with blue epoxy resin to define porosity. (a) D1 inclusion-rich dolomite developing along a

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Locally, D2 dolomites are difficult to distinguish from D1, although it appears likely that D2 dolomites are in part centred over D1 cores (Fig. 32a, b). This observation could be interpreted as evidence that D1 and D2 are in fact related to the same stage of dolomitization. Petrographic analysis shows that the D2 phase varies from anhedral to euhedral dependent on the degree to which it is replacive and also the packing of crystals (‘overdolomitization’ – that is, all calcite replaced and porosity occluded). The dolomites are locally mimetic (e.g. of radiolitid rudists). Intercrystalline and mouldic (after Orbitolina sp.) porosity are evident (Fig. 31c, d). In CL, D2 dolomites are either bright red/oranges with reddish cores, or dull –milky purple colours. Petrographic, fluid inclusion and organic matter maturity data studies indicate burial depths of c. 2.5 km (80 8C) and that D2 dolomitization occurred at that depth, implying that D2 dolomites can not be defined as hydrothermal sensu-stricto (Lapponi et al. 2010). Burial to 2.5 km is thought to have occurred during the middle Miocene (Homke et al. 2004). An important phase of hydraulic fracturing (Fig. 32c), corrosion and vug development (Fig. 31f –h), affecting the D2 dolomites, precedes precipitation of D3–D5 dolomites. This stage (actually several stages) is easily identifiable in the field and is confirmed in thin-section. D3–D5 dolomites are characterized by white saddle dolomites with undulose extinction and well developed curved crystal facies. D3–D5 dolomites fill vugs, fractures and moulds, and are often coarsely crystalline. In cavern and vug filling saddle dolomite botryoids developed adjacent to the normal fault in Figure 23 up to nine phases of saddle dolomite and iron sulphide precipitation can be identified. In thin-section however, only three phases of saddle dolomite development can be systematically distinguished. Petrographically, D3– D5 dolomites are very distinctive. Crystal terminations are usually euhedral, and in CL the dolomites are distinctively zoned bright orange/red and reddish brown

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colours (Fig. 32d, e). D3 –D5 dolomites are commonly pore/fracture-filling and overgrow dolomite rhombs of D2 origin. Several phases of fracturing (including crackle or exploded texture fractures) are also associated with these dolomites, with a complex and intricate relationship between the fracturing and dolomite phases. Temperatures of homogenization in primary fluid inclusions in zoned saddle dolomites (D3 –D5) and late calcites (LC) record temperatures between 80 –130 8C and the development of hydrothermal pulses during precipitation (Figs 30 & 33). Oil inclusions are also evident. These data prove a hydrothermal origin (i.e. the introduced fluids responsible for saddle dolomites were hotter than the surrounding country rock). Organic matter studies adjacent to faults confirm localized elevated temperatures. Several phases of calcite cementation postdate the saddle dolomites. Late calcites phases LC1 –LC5 are very distinctive and well-zoned in thin-section. The full zonal sequence starts from non-luminescent cement and passes gradually into bright luminescent cement, then to sector-zoned cements. Thin-sections show evidence for these calcites being fracture-fed (Fig. 32f, g). The late calcite cements show a trend from initially depleted (low) to enriched (higher) d13C values and a progressive increase of d18O (Fig. 30, Lapponi et al. 2010). The ‘switch’ to calcite cementation indicates a probable change of thermodynamic conditions and/or hydrothermal pore waters, becoming less magnesium and more calcitic (alkaline). The fine zonation observed in CL suggests micro-variations in chemistry of pore fluids although the precise control on these variations is uncertain (changes in original input fluid compositions or effects of rock-buffering as fluids cemented the rock). Fluid inclusions range from 120–130 8C indicating continued hydrothermal conditions during late calcite precipitation (Fig. 33). A suite of associated minerals, including fluorite, kaolinite, barite and dickite are also associated with the late calcites. LC1–LC5 incompletely fill vugs and fractures.

Fig. 31. (Continued) stylolite within pelloidal dasycladacaen-rich platform-top micritic facies. D1 dolomite overprints early calcite cements. (b) Scattered inclusion-rich D1 dolomite rhombs developed in a Chondrodonta storm bed (see also Fig. 17). (c) Tight (nonporous), undolomitized orbitolina packstone facies with early marine calcite cements. (d) Totally dolomitized (D2) porous Orbitolina packestone facies. The Orbitolina are dissolved out to form well developed mouldic porosity. Facies in C can be walked into facies in D over a few metres. Burial dolomitization thus has a dramatic effect on porosity/reservoir creation. (e) D2 dolomites, with inclusion rich centres and clear zoned outers. Saddle and late calcite cements partially fill the intercrystalline pore space. (f) Dissolution of D2 dolomites and partial cementation by D3 –D5 saddle dolomites. (g) D2 dolomites showing pervasive dissolution and vug development prior to precipitation of vug-lining and pore-filling D3 –D5 saddle dolomite cements. Geopetal fill of uplift related pink vadose silts also evident (arrow). (h) Saddle dolomite partially lining vug in D2 dolomite. Saddle dolomite and late calcite filled veins are also evident. Note geopetal fill prior to precipitation of saddle dolomite. Compare to Figure 27 (same location).

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Fig. 32. (a & b) Plain Polarized light (PPL) and cathodluminescence (CL) pair of photographs highlighting burial D1 and later D2–D4 dolomites. Cores of the dolomite are D1 phase, consisting of bright reddish euhedral rhombs. These are etched and overgrown by D2 phase milky then dull purple dolomite with ragged/etched outlines, followed by brown euhedral D3 dolomite. The final dolomite phase is a well-zoned (thin bright orange fading to red then dull with a

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Fig. 33. Summary of burial dolomite and late calcite (LC) fluid inclusion data. The saddle dolomites have generally higher temperatures than mosaic dolomites, and a distinct cooling trend is evident to LC1 calcites associated with a decrease in salinity/influx of low salinity water. Occurrence of hydrocarbon inclusions also plotted at top.

Uplift-related diagenesis

Reservoir characterization

The last diagenetic phases identified appear to be related to uplift, cooling and exposure of the Anaran Anticline. These include yellow luminescent vadose silts (typically developed in fracture and vug cores), dissolution and dedolomitization and calcrete formation with local dripstones/speleothem textures. Homke et al. (2004) interpreted the late uplift of the Anaran Anticline to be .3 km with respect to the foreland.

Outcrop data record the development of complex, well connected, three-dimensional dolomite reservoir bodies controlled by fractures (joints and faults), facies/stratal architecture and diagenetic heterogeneity. Dolomitization caused porosity redistribution (enlargement, destruction, formation of new porosity/pore types). Despite apparent complexity, we have shown that an understanding of which fault/joint systems were active during

Fig. 32. (Continued) final thin red/orange subzone) D4 phase. The subsequent vug-fill phase consists of a streaky zoned brown/dull calcite cement (LC5). (c) Whole thin-section scan highlighting nature of hydraulic fracturing and zebra fabrics. Bands of dark, mimetically-dolomitized matrix sediment are separated by bands of saddle dolomite. Remnant vuggy pores in the banded microfacies are partially infilled by coarse sparry calcite. There are minor late microfractures which cross-cut the dolomite. (d & e) PPL-CL pair highlighting hydrothermal dolomite and late calcite phases. Initial dolomitization (D2) is inclusion-rich subhedral-anhedral crystals which replace matrix. These dolomites have a uniform dull dark brown character in CL. A later pore-filling dolomite is then precipitated (D3). The rhombs are euhedral and grow into open pore-space. These dolomites often form the nucleus for later, D4 pore-filling dolomites. The pores formed initially as a result of a dissolution event where bioclasts are preferentially removed. D4 euhedral dolomite crystals continue to infill porosity. They have more limpid properties in PPL (compared to the earlier dolomite phases), yet are zoned. In CL, they display concentric multiple zones of oranges and reds. The final dolomite phase is D5 with a dull orange luminescence. Etched LC5 calcite infills remaining pore space and has a yellow-brown CL character. (f & g) PPL-CL pair highlighting late calcite cements. D1 dolomite crystals show a brownish followed by red luminescence and euhedral zonation, whilst D2 dolomite has a darker brownish luminescence. A brecciation/fracture event then develops porosity which is filled with LC2 dull brown to non-luminescent concentrically zoned/subzoned calcite spar, followed by LC3 non-luminescent to dull brown concentrically zoned/subzoned calcite spar, and finally by LC4 moderate brown to bright yellow luminescent concentrically zoned/subzoned calcite.

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dolomitization, coupled with systematic mapping of dolomite distribution within a facies and sequence stratigraphic framework allows establishment of a qualitative predictive template (Figs 10, 28 & 29). This template can be used to aid reservoir characterization and model building in the subsurface. In order to assist characterization of fracture controlled dolomite reservoirs in the subsurface, an outcrop reservoir model was built using the NW Dome outcrops (Sharp et al. 2006). This model was built up in three layers to capture the different types of reservoir heterogeneity (depositional, diagenetic and fracture, Fig. 34). Building the first (depositional) layer used a vertical synthetic well (outcrop section) to establish zonation, primary facies, microfacies, poro-perm data, and depositional architecture. Depositional geometry, facies and layering was built away from the control section using geometries observed at outcrop. The depositional model was populated with poro-perm data collected at outcrop and using the methodology of Lønøy (2006). The diagenetic (dolomite, karst) and fracture (faults, joints) layers were then superimposed on the depositional layers. It is beyond the scope of this paper to describe the full model (see Sharp et al. 2006 for more details). However, it is relevant to describe how the fracture controlled dolomites were modelled, and the main challenges in reservoir characterization and modelling of fracture controlled dolomites. Modelling of the dolomite reservoirs involved detailed photo interpretation of the distribution and geometry of dolomite bodies observed at outcrop in dip, strike and plan sections. These interpretations were used to build dolomite ‘geobodies’ (objects) which formed a unique reservoir layer that replaced pre-existing limestone matrix in the reservoir model. Within the objects vertical and lateral trends were input for percentage limestone versus dolomite. Vertical and lateral poro-perm trends were input in to the dolomitized parts of the geobodies based on outcrop poro-perm data. Facies belts were also used to model dolomite bodies with gradational vertical and lateral transitions to limestone. The average fraction of dolomite was given manually for each layer and vertical and lateral trends input. Both methods produced realistic dolomite geobodies which recreated outcrop geometries (Fig. 34). A third type of dolomite modelled was laterally extensive stratabound dolomite which occurred at the same stratigraphic level throughout the model area. These dolomites were conditioned to synthetic well data (outcrop sections) and allowed to fill the entire model laterally. Vuggy and cavernous porosity is associated with D2 and D3 –D5 dolomites. Vugs were observed to

be in the order of 5 mm to 10 cm across at outcrop. Several vug types were identified: (a) Non-fabric selective, random. (b) Non-fabric selective, occurring towards the tops of depositional cycles, with cycle tops marked by thin mudstone/marl permeability barriers. (c) Stratabound elongate vugs associated with zebra saddle dolomite cements spatially related to fault zones. Vuggy porosity estimates based on image analysis of outcrop photographs gave values of up to 20%. However, 90% of these vugs had no cement lining, and were thus interpreted to be surficial weathering features. Vuggy porosity with cement lining is in the order of 5–10%. These cement-lined vugs were interpreted as representative of vuggy porosity that would be present in the subsurface. The percentage of vuggy porosity was added to dolomite matrix values on a zone by zone basis. Lateral trends were given a wide variability, but an increase in vuggy porosity towards fault zones and towards the tops of depositional cycles was implemented to reflect outcrop observations. Similarly, cavernous porosity was only observed and hence modelled adjacent to fault zones, and formed less that 5% of the total rock volume in the reservoir model. Modelling of the D2 and D3– D5 dolomites as objects and facies trends worked well for the outcrop model, as constraints on the extent and dimensions are easily obtained. However, predicting the length-scale distribution of porositypermeability variations within dolomite bodies is more challenging, requiring time consuming lateral and vertical sampling in the field within individual dolomite tongues developed within individual primary facies types and individual fault zones. Our preliminary poro-perm data collected at outcrop indicate a relatively uniform log-normal distribution of porosity and permeability for crystalline D2 dolomites regardless of facies (Sharp et al. 2006), but this is based on a relatively limited sample suit (250 samples) and is not considered adequate to capture the true variability of dolomite porosity and permeability. Further systematic sampling is thus required on a facies by facies basis in order to develop a pre- and postdolomite poro-perm dataset. Similarly, systematic vertical and lateral sampling transects are required in the footwall and hanging wall to normal faults for individual facies types to capture fault zone heterogeneity. This work is time consuming, but is key to quantitative and realistic representation of reservoir properties. With regard to subsurface reservoir characterization, the established predictive template, quantitative dataset and outcrop model goes a

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Fig. 34. Workflow for modelling and reservoir characterization of dolomite bodies using the Anaran outcrops. (a & b) A 1.2  1.6 km  350 m thick area of outcrop was modelled (Mauddud Mbr þ Lower Sarvak Fm). The model was built up in three layers (depositional, diagenetic, fracture). The depositional layer (c) used a vertical synthetic well (outcrop section) to establish zonation, facies/microfacies and depositional architecture. The depositional model was populated with ‘facies’ poro-perm data collected at outcrop (d). The diagenetic bodies (dolomites) were then superimposed as geobody objects (plumes, sheets, patchy sheets) based on outcrop geometries (e), and dolomite porosity/permeability values replaced ‘facies’ values (f). Fracture (faults, joints) porosity and permeability layers were then superimposed on both depositional and diagenetic layers (g). The full model building workflow is described in Sharp et al. (2006).

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considerable way to aid subsurface model building, and in particular allows ‘weighting’ of which primary facies and faults are more susceptible to dolomitization. However, subsurface data in the area is limited to widely spaced wells and 2D seismic data, resulting in considerable uncertainty with regard to property modelling. Forward stratigraphic modelling conditioned to well, seismic and outcrop data has thus been used to establish the primary (depositional) heterogeneity (Fig. 35, Embry et al. 2008). Stochastic modelling of dolomite bodies has been superimposed on this depositional template. The stochastic modelling is based on the field observations of dolomite occurrence along faults and as stratabound bodies following specific horizons (e.g. permeability pathways such as karst and ‘grainy’ carbonate facies, and beneath permeability barriers). A ‘greater susceptibility to dolomitization’ weighting was given to faults observed on the seismic with the same trends as faults know to be associated with dolomitization at outcrop. Similarly, specific facies types and horizons were given a ‘susceptibility to dolomitization’ weighting based on the outcrop framework. In addition, seismic forward modelling of the outcrop sections was undertaken to produce synthetic seismic sections which were used for seismic facies classification and to aid identification of dolomite bodies on 2D subsurface seismic. This combined methodology allowed a realistic representation of the external form of the dolomite reservoirs (Fig. 35). However, the results still fall short of what is needed to fully characterize and accurately model the dolomite reservoirs in the subsurface. Two areas in particular remain a challenge: (a) Prediction of lateral and vertical extent of individual dolomitic bodies away from feeder points (fault zones). (b) Vertical and horizontal length scales of variability of porosity and permeability within dolomitic bodies. Published and unpublished data from hydrothermal dolomite reservoirs have indicated that well-tied multi-attribute analysis of 3D seismic is the key to accurate imaging of the external form of hydrothermal dolomite bodies (Sagan & Hart 2006; Davies & Smith 2006). Where 3D data is lacking, reactive transport modelling may be used to directly model the development of dolomite plumes, and to develop an understanding of the controls on dolomite body size and orientation (e.g. Yao & Demicco 1995, 1997; Whitaker et al. 2004; Jones & Xiao 2005). Predicting the vertical and horizontal length scales of variability of dolomite porosity and permeability can only be achieved by the establishment of a facies-based quantitative outcrop and subsurface (core) database.

Predicting the distribution and geometry of chalky/microporous limestones, both at outcrop and in the subsurface, is also a major challenge. Geometrical data is more difficult to gain from the outcrop data as these lithologies do not weather as clearly visible bodies as the dolomites do. Obtaining the dimensions of these lithologies will thus involve more systematic outcrop sampling and mapping. To date, the majority of our outcrop data point to a clear primary facies control on the occurrence of chalky and microporous limestones (typically within platform margin to interior rudistic and shoal facies), and length scales well beyond the scale of the outcrop model and dolomite body distribution.

Regional comparisons and discussion In order to establish how representative the data from the Anaran study area are, it is useful to compare to other known fracture-controlled dolomite occurrences. It is also informative to compare fracture-fed dolomites developed in compressive (Zagros), extensional (Gulf of Suez) and transtensional (North America) settings.

Middle East – Zagros fold belt and Arabian plate Outcrop study of other anticlines within the Zagros fold belt has revealed that fracture-controlled dolomites are common. In the Lurestan area of Iran alone, fracture-fed and associated stratabound dolomites have been observed in the Khami and Bangestan groups in the following areas; Kuh-E-Anjir, Kabir Kuh, Kuh-E-Safid, Kuh-EChenareh, Siah Kuh, Kuh-E-Rit, Kuh-E-Kurnas, Kuh-E-Sarkan, Kuh-E-Sultan, and in numerous anticlines exposed along the Tang-E-Haft river section (Fig. 36a). In the majority of these outcrops, the dolomites are developed along normal faults (e.g. spectacularly exposed in Kuh-E-Anjir) and show a strong facies/depositional architecture control on lateral extent (e.g. Kuh-E-Safid, Fig. 36c). These observations are directly comparable to data from Anaran. In contrast, in the Chenareh, Rit and Kurnas anticlines, located in the more internal parts of the Zagros Simply Folded belt, dolomite is observed developed along Cenozoic thrust faults which cut neritic facies of the Sarvak Fm (Fig. 36b). The thrust faults are associated with chaotic, mosaic and crackle breccias, composed of floating clasts of ferroan dolomite (cf. D2 in Anaran) cemented by white saddle dolomite (cf. D3–D5 in Anaran). Layer parallel and shear zebra fabrics cemented by white saddle dolomite are also well developed, and can be mapped short distances

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Fig. 35. Reservoir modelling in the subsurface. Forward stratigraphic modelling of depositional architecture conditioned to wells and outcrop logs can be used to build primary depositional facies as input to reservoir models (a). Seismic modelling (b) of outcrop geometries and object based models of plug and stratabound dolomites (c) can then be used to superimpose diagenetic bodies, which can be combined to produce a final matrix reservoir model integrating primary and secondary reservoir facies (d).

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Fig. 36. (a) Known occurrences (outcrop and subsurface) of fracture-controlled dolomites in the Lurestan region of SW Iran. (b) D2 and D3–D5 dolomite developed in hanging wall of reverse fault (arrowed) developed on NE limb of Kurnas Anticline, Lurestan. Also note laterally extensive stratabound dolomites in footwall region, developed within the Sarvak Formation. (c) Stratabound dolomites developed within the Khami Group, Kuh-E-Safid Anticline near Khoram-abad, Lurestan. Dolomites occur beneath low permeability shelf margin clinoforms and in platform top facies. Note absence of dolomite in ‘grainy’ margin facies. Image courtesy of IFP (Frans van Buchem). (d) Outcrop photograph of dolomite body developed along normal fault within Precambrian carbonates. Jebel Akhdar, Oman.

(10–50 m) away from the thrust faults. They are developed in both the footwall and hanging wall, although volumetrically are more significant in the hanging wall. Laterally more extensive (km scale) stratabound ferroan dolomite bodies (cf. D2 in Anaran) can also be mapped away from the thrust faults (e.g. Rit, Kurnas – Fig. 36b), and occur either beneath aquitards and/or following precursor D1 dolomitic limestones (slope and platform interior facies). These dolomites are volumetrically most significant adjacent to the thrust faults, and progressively die out away from the fault. They appear to be more prevalent in the hanging wall (e.g. Kurnas Anticline – Fig. 36b). Exposures of the basinal equivalent of the Sarvak Fm (Garau Fm) in TangE-Haft reveal that fracture-fed and stratabound dolomites are also common in these facies, and thus not confined to platformal settings (cf. Goff 2005). In the Fars province of southern Iran, fracture-fed and stratabound dolomites in the Khormuj, Khartang and Gach anticlines have affected the interval from

top Jurassic to Upper Eocene. In the Khartang Anticline dolomites within the Surmeh, Hith and Fahliyan formations are concentrated along normal conjugate faults which feed stratabound bodies capped by aquitards (in this case evaporites and cycle capping mudstones). The normal faults are located on the steeper dipping SW limb of the anticline. Importantly, dolomite bodies in the Fars province are preferentially development along N–S/NE– SW trending basement lineaments. In the Zagros Mountains of Iraq, dolomitization of the Lower and Mid Cretaceous Qamchuqa Group is widespread (Al Shadidi et al. 1995). A fracture origin is postulated based on proximity to the Iranian examples described in this paper. Also in Iraq, Goff (2005) described fracture-controlled dolomitization of the Late Jurassic Surmeh Fm, with dolomitization focused along the Surmeh platform margin. At outcrop in Oman, fracturecontrolled dolomites are well exposed cutting preCambrian carbonates in Jebel Akhdar (Fig. 36d)

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and also cut Cretaceous limestones in the Adams Foothills, where they are volumetrically subordinate to ‘leached/chalky’ microporous limestones. In the Ghawar field of Saudi Arabia, Cantrell et al. (2004) attribute dolomites in the Arab D Formation to a fracture origin.

Middle East – Gulf of Suez Fracture-controlled hydrothermal dolomites and calcites have been described from crustal-scale (25 km wide, 50 km long) tilted fault blocks of the Suez Rift (e.g. Clegg et al. 1998). The dolomites described by Clegg et al. (1998) are spatially limited and volumetrically small by comparison to the Iranian examples, although the dolomitization mechanism invoked is comparable. Outcrops in the footwall to the Gebel Araba and Hammam Faraun fault blocks however, located on the eastern (Sinai) side of the Suez Rift (Patton et al. 1994; Moustafa & Abdeen 1992; Moustafa 1996), expose areally extensive fracture-fed and stratabound dolomites (Fig. 37a–c). Dolomites within the Hammam Faraun Fault block have three distinct occurrences: (a) As steep dipping to vertical discontinuous sheets along major normal fault zones (for example the Hammam Faraun Fault, which has over 5 km of normal displacement, Fig. 37a, b). Dolomitization is typically of the fault breccias and adjacent country rock to produce dolomitized ‘halos’ (cf. Davies & Smith 2006) which extend short distances (,50 m) into the hanging wall and footwall. (b) Dolomites adjacent to and above rift-related Oligocene volcanic dykes and sills, again forming localized ‘halos’ (1–10 m wide). Dolomitized joints can be mapped away from the edges of dykes and sills. The dykes are often injected along fault planes. These dolomites are comparable to those described by Nader et al. (2007) associated with volcanics in the Jurassic of Lebanon. (c) As laterally extensive (1–4 km) stratabound bodies, especially within ‘grainier’ carbonate facies. This situation is well exposed in coastal exposures of the Eocene-aged Thebes Fm in the immediate footwall of the Hammam Faraun Fault (Fig. 37b), where sub-vertical dolomite bodies developed along the Hammam Faraun Fault can be traced into stratabound dolomites which preferentially develop within slump sheets, debris flows and grain-benthic foram rich turbidites. Bioclasts (corals, nummulites) in these facies are often dissolved to form moulds and vugs which are lined by saddle dolomite crystals

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(cf. Orbitolina in Iran, Fig. 31c, d). The stratabound dolomites are locally associated with a well developed zebra shear fabric. The zebra fabrics are filled by saddle dolomite and show a consistent sense of shear away from the fault zone. In sections parallel to the strike of the fault, the zebra fabrics are seen to be domal in nature, and associated with bilateral shear microfractures (ladders) and rimmed microfracture fabrics (Fig. 37c, Davies & Smith 2006). These textures are directly comparable to those observed adjacent to normal faults in Iran (this paper) and to those described from Canada (Fig. 37f, cf. Figs 5– 15, Davies 2004; Davies & Smith 2006; Vandeginste et al. 2005). The fracture-controlled dolomites of the Hammam Faraun Fault zone are also associated with chalky/ microporous limestone halos, and individual stratabound dolomites can be ‘walked out’ into microporous limestones. Contact metamorphism and baking of the limestone country rock adjacent to volcanic dykes, sills and mineralized joints is also present. In one example at Hammam Faraun hot springs, the Late Cretaceous Sudr Fm, which forms one of the major source rock intervals in the region (Patton et al. 1994), shows localized baking, bitumen formation and contact metamorphism between two sub-vertical fractures (joints). This contact metamorphism is lacking when the Sudr Fm is walked out away from the fractures. This situation is comparable to that observed within organic rich mudstone units adjacent to major normal faults in Anaran, Iran. The present day location of hot springs (Hammam Faraun translates as Pharaohs baths), oil/bitumen seeps, karstic caverns, active normal faulting, and dolomite bodies along the Hammam Faraun Fault and other block-bounding faults (e.g. Gebel Araba, El Tor and Hammam Musa) in the Suez Rift indicate that hydrothermal systems (and diagenesis?) are active present day. The Sinai outcrops may thus form a unique area where fracture-controlled diagenesis can be studied as a modern active process.

Western Canada and NE USA Comparisons between the Iranian data and well documented Cambrian, Devonian and Mississippian hydrothermal/high temperature dolomite (HTD) reservoirs of western Canada and the north eastern USA reveal marked similarities with regard external form (Fig. 37d, e), textural relationships, paragenesis and models of structural emplacement (e.g. Yao & Demicco 1995, 1997; Swennen et al. 2003; Davies 2004; Smith 2006; Davies & Smith 2006; Vandeginste et al. 2005). However, several

390 I. SHARP ET AL. Fig. 37. (a) View of the Hammam Faraun fault scarp, Gulf of Suez. Footwall on right, hanging wall on left. Fault displacement is 5 km. Fracture-controlled dolomite bodies are well developed along the fault and extend into the footwall region (Fig. 37 b). (b) Detail of dolomites in the immediate footwall. The dolomites can be mapped up the fault and extend into the footwall as stratabound sheets following bioclast and grain-rich debrites and turbidites. Trace of fault indicated (c) Domal saddle dolomite shear fabric developed adjacent to the fault zone. Note bilateral shear microfracture defining left-hand edge of dome fabric. Shear is into the page. (d, e & f). Fracture-fed dolomites from the Cambrian of the Canadian Rockies. (d) Helicopter photograph of spectacular vertical dolomite pipe/plume (up to 200 m high – arrowed) developed in outcrops above Lake O’Hare. The external geometry is directly comparable to plumes developed in Iran. (e) Dolomite plumes developed along joints and as stratabound bodies. Note ‘bulbous’ lower boundaries and stratabound top. Outcrop is approximately 50 m high. Yoho Glacier region, Canadian Rockies. Arrows indicate dolomite bodies. (f) Domal and inclined zebra shear fabrics. Sense of shear is away from viewer in lower image, and towards the left in upper image. Both outcrops are c. 50 cm high. These fabrics are directly comparable to those developed in Iran and Gulf of Suez.

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notable differences occur, which are worthy of further discussion: (a) The importance of joints during the dolomitization process in the Iranian outcrop examples. There is limited published data to date addressing the relationship between Canadian/North American HTD reservoirs and jointing, although Davies & Wendte (2005, their Figs 39 & 40) illustrate abrupt dolomitelimestone boundaries along shear microfractures. Focus has mainly been on faulting and stratal controls. The Iranian data has shown that closely spaced stratabound joints provided one of the main pathways by which dolomitizing fluids were able to access and dolomitize large areas of low permeability country rock. This includes dolomitization of argillaceous limestones which otherwise will behave like a cap rock to dolomitizing fluids. Based on our data from Iran we feel this point justifies further study in the Canadian/North American examples, as understanding the relationship between joint density and facies could be key when trying to characterize HTD reservoirs in the subsurface, and also when trying to ascertain the age of emplacement of the North American HTD. (b) Late stage of emplacement as part of Zagros folding for Iranian examples as opposed to an early, shallow emplacement model for North American examples. In the Zagros the late stage of emplacement, and an intimate relationship to development of the Mesopotamian foreland basin and Zagros fold belt appears clear. HTD dolomite is mapped as developing along joints and normal and reverse faults which are intimately related to folds dated as developing between 8 to 5 Ma. In contrast, the Cambrian, Devonian and Mississippian HTD’s of western Canada and North America are interpreted to have formed either syn-depositionally or shortly after burial at relatively shallow depths, and certainly prior to Laramide deformation (Yao & Demicco 1995, 1997; Swennen et al. 2003; Davies & Wendte 2005; Davies & Smith 2006). (c) Strike slip v. extensional and thrust faulting. An association of the HTD to thrust as well as normal faults is clear in the Iranian examples. In contrast, in the Canadian and North American examples, an association with wrench/ strike slip faulting is well established, and locally a relationship to extensional faults (e.g. Mississippian Debolt Fm in NE British Columbia, Davies & Berger 2004). No clear relationship to thrusts/reverse faults has been established.

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(d) The development of subtle domes (Iranian examples) as opposed to sags (Canadian– North American examples) above dolomite plumes and faults. The identification of subtle doming and associated fracturing above dolomite plumes in the Iranian examples is quite different from the sags developed in the Canadian and North American examples (Davies & Smith 2006; Sagen & Hart 2006). Possibly this is related to an overall addition of material during the dolomitization process, or ‘fossilization’ of the Iranian dolomite plumes before subsequent fluid withdrawal, hydrostatic pressure drop and/or corrosion and collapse? It may also relate to transtensional tectonics being dominant in the Canadian– North American examples, resulting in greater space development and collapse. In summary, the similarity in external form, textural development and paragenesis of global fracturecontrolled dolomites is striking. This implies a uniform emplacement and dolomitization mechanism for fracture-controlled dolomite, regardless of age and to some degree structural setting. All examples emphasize the importance of the pre-existing stratigraphic architecture on the resultant dolomite body distribution and external form. All examples also emphasize the importance of faults, and in particular hydraulic fracturing during co seismic fault rupture to allow advancement of dolomitizing fluids into the country rock. Not all examples address the importance of joints (density, orientation, spacing, aperture) as a mechanism for dolomitizing large areas of country rock. We feel a major lesson learnt from the Iranian case study is the importance of establishing a mechanical (joint) stratigraphy to predict dolomite distribution. This may have been underestimated in the other case studies. What perhaps is most surprising is the textural and paragenetic similarity developed along normal, reverse and strike-slip faults in either compressional (Zagros), extensional (Suez) or transtensional (Canada–NE USA) regimes. Fault breccias, bedding parallel and inclined shear zebra fabrics are remarkably similar in all examples studied regardless of structural setting. This appears to imply a similar dolomitization and hydraulic fracturation process regardless of structural setting. In two of the discussed case studies (Anaran and Gulf of Suez) there is evidence to support the theory of local ‘forced maturation’ of organic rich formations along faults and fractures associated with focused elevated heat flow, as originally suggested by Davies (2004, his Fig. 16) and Esteban & Taberner (2003). We feel a detailed study of this phenomenon would be of great benefit to the debate on the presence or absence of pulsed

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hydrothermal systems and dolomites developed along faults. Systematic and detailed lateral sampling of units affected by such contact metamorphism would give a good understanding of temperature gradients associated with fluid flow along faults, and effectively end the hydrothermal dolomite debate (Machel & Lonnee 2002). The relationship between late microporosity creation in leached/chalky limestones and HTD development remains somewhat enigmatic, but there is a clear relationship locally (e.g. Ladyfern field, Davies & Smith 2006). Further work is needed to understand this relationship. From a process standpoint, many questions remain, including the source and volume of the dolomitizing fluids, relationship to hydrocarbon migration (if any), and the process of dolomite front advance and fluid flow within the country rocks and fractures (advection, diffusion, convection). Based on our Iranian data, we propose three models of dolomitization; (a) Focus of massive hydrothermal dolomites along the palaeo-shelf margin, with a transition to stratabound dolomites, dolomitic limestones and ‘leached/chalky’ microporous limestones in the shelf interior. In this model, the source of Mg might be basinal shales of the Garau Basin which fringe the shelf margin in Lurestan, or underlying Jurassic and Triassic evaporites. This model is comparable to that suggested for the Cambrian and Devonian of Canada (Yao & Demicco 1995, 1997; Davies & Smith 2006) and the Jurassic of Iran –Iraq (Goff 2005). (b) A deep-seated basement fault control, with elevated heat flow along long-lived crustal scale N –S/NE –SW trending faults. (c) A combination of both of the above. On balance, we favour model ‘b’, as dolomite bodies developed in the Fars province of SW Iran are far removed from the shelf margin and show a preferential development along long-lived basement lineaments associated with elevated heat flow.

Conclusions A revised lithostratigraphic, biostratigraphic and sequence stratigraphic template is presented for the Upper Khami and Bangestan Groups in Lurestan, SW Iran. Reservoir, source, seal, structure and diagenesis are incorporated into this framework, and form a powerful reservoir characterization tool with which to address subsurface and outcrop datasets. The importance of tectonics on sequence architecture and karst development in the study area is stressed, in particular during the Aptian (Intra Dariyan), Latest Albian (Top Mauddud) and latest Cenomanian-Turonian (Upper Sarvak). Outcrop

data indicate that the Mauddud Member should be assigned to the Kazhdumi Formation, and not to the overlying Sarvak Formation. Outcrop data have also been utilized to demonstrate stratigraphic and structural controls on the development of regional scale dolomite bodies. Dolomite bodies are; (a) Fracture-fed vertically, utilizing faults and joints to dolomitize large areas of relatively low permeability country rock. (b) Stratabound laterally, following ‘permeability pathways’ (e.g. Karst) and/or beneath aquitards. (c) Associated with both diagenetic (dolomite front) and fracture-defined edges (joints and faults). Saddle dolomites and associated saddle dolomitecemented shear fabrics (zebras) and floating clast jig-saw breccias are spatially related to fault zones. They are interpreted to have developed due to pulsed pressure release and hydraulic brecciation during co-seismic deformation (fault rupture). Pulsed fault rupture results in multiple corrosion/ cementation phases. Field relationships indicate that D2– D5 dolomites formed slightly pre- to synZagros folding (between 8 –5 Ma), whist fluid inclusion and organic matter maturity data indicate that D2 dolomitization occurred at 80 8C, which equates to 2.5 km burial depths. Temperatures of homogenization in primary fluid inclusions in zoned saddle dolomites (D3–D5) and late calcites (LC) record temperatures between 80 –130 8C, thus proving the existence of a hydrothermal system (i.e. introduced fluids were hotter than the surrounding country rock). These data and observations are in good agreement with the observations and models of Yao & Demicco (1995, 1997), Swennen et al. (2003) and Davies & Smith (2006) for the generation and emplacement of hydrothermal dolomite bodies. Comparison to published examples and our own data from elsewhere in the Middle East, indicate that fracture-controlled diagenesis and dolomitization is more widespread than previously thought, and indeed may be the predominant dolomitization mechanism in a number of prolific Middle East reservoirs. Our data also indicate a common hydraulic fracturing process is applicable along normal, reverse and strike slip faults in either compressional (Zagros), extensional (Suez) or transtensional (Canada, North America) regimes. The importance of an integrated structural, sedimentological, stratigraphic, petrographic and geochemical approach to understanding fracturecontrolled dolomites and diagenesis is stressed. In this paper, we emphasize the importance of detailed field observations, and in particular of systematic

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documentation of cross-cutting relationships at the outcrop scale, especially of joint systems. Petrographic and geochemical studies in isolation are not adequate to fully describe fracture-controlled diagenesis, as they lack the spatial variability that can be systematically documented at outcrop. They are crucial however, to constrain background and emplacement temperatures, fluid composition and source, and establishment of a detailed cement stratigraphy. From a reservoir characterization perspective dolomitization caused porosity redistribution, enlargement, and formation of new porosity. The benefit of outcrop study is clear, allowing unequivocal documentation of the relationship between structure, stratigraphy and diagenesis. These data can be used as a qualitative and semi-quantitative predictive tool and framework for modelling dolomite distribution in the subsurface. However, much remains to be done on reservoir characterization, in particular with regard to the establishment of quantitative datasets. Key questions include; (a) Is there a relationship between fault displacement, fault damage zone width, and dolomite body size? In particular, is there a relationship between the distance stratabound dolomites extend away from individual faults and fault displacement? (b) Is there a systematic variation in the length scales of dolomite porosity and permeability variability within specific primary facies? (c) At what distance from the fracture feeder is secondary porosity creation optimal, and under what conditions does ‘over-dolomitization’ lead to the destruction of secondary porosity? (d) What is the spatial relationship between fracture-fed dolomites and ‘leached/chalky’ microporous limestones, over what distances do leached limestones develop, and what is the process that creates ‘leached/chalky’ limestones. Our outcrop study has gone someway to address these questions, particularly with regard to the stratal and fracture control and resultant geometries, but it is clear that further detailed, systematic field, petrographic and modelling work on dolomite bodies is needed to obtain more quantitative answers. This work has been carried out as part of a joint study between Norsk Hydro (now Statoil) and NIOC (National Iranian Oil Company). The manuscript greatly benefitted from editorial input by Graham Davies, Mateu Esteban and Frans van Buchem. Stephen Packer and Esmeralda Caus are acknowledged for dating benthic foraminifera and associated fauna. Statoil and NIOC are thanked for permission to publish these data. Stian Soltvedt, Jon Ineson and Jean-Christophe Embry are acknowledged for contribution and discussion concerning reservoir characterization.

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Contrasting fluid events giving rise to apparently similar diagenetic products; late-stage dolomite cements from the Southern Alps and central Apennines, Italy P. RONCHI1, A. DI GIULIO2, A. CERIANI2* & P. SCOTTI1 1

Eni S.p.A. Exploration and Production Division, Via Emilia 1, 20097 S. Donato Milanese, Italy 2

Dipartimento di Scienze della Terra, Universita` di Pavia, Via Ferrata 1, 27100 Pavia, Italy *Corresponding author (e-mail: [email protected]) Abstract: Precipitation of dolomite cements in Jurassic carbonate platform sediments and slope breccias has been studied from well cores and outcrops of the central Southern Alps and central Apennines in Italy. In both areas, an initial, massive dolomite replacement was followed by multiphase precipitation of dolomite cements. The replacement occurred during burial, in a passive margin regime, in response to compaction-driven flow of formational fluids. This interpretation is based on results from fluid inclusion and stable isotope analyses which have been related to the thermal history. The dolomite cements precipitated when both areas were involved in collisional tectonics. In spite of the similar diagenetic evolution, the fluids causing dolomite cementation in each case were compositionally different. In the Alps a decrease in salinity was recorded from sea water to brackish fluids, whereas in the Apennines an increase in salinity from sea water up to .10% NaCl equivalent was recorded. The remarkable salinity differences in diagenetic fluids are considered to be related to the different sub-aerial relief of the two belts during dolomite precipitation. In the Alps, the dilution of fluids is related to the infiltration of meteoric waters from the mountain chain, that was widely emergent. In the Apennines, dolomite cements precipitated whilst the structural units were still widely submerged, preventing meteoric dilution of cementing fluids and promoting an increase in salinity through mixing with fluids rising from older evaporate-bearing layers. In both Alpine and Apennine cases, the same diagenetic trend is observed in thrust-fold belt and foreland basin units; in both structural systems the diagenetic events start precipitating dolomite cements in the inner part of the collision zone and then the diagenetic processes migrate towards the foreland basin along with the structural evolution of the area.

In carbonate petroleum systems, dolostones commonly show higher porosity and permeability with respect to their limestone counterparts. For this reason, dolomitized bodies often represent an exploration target for the oil industry. Moreover, an understanding of the dolomitization processes is helpful for predicting reservoir geometry, continuity and petrophysical characteristics and therefore it is fundamental for reservoir development. Many models have been proposed for massive dolomitization (see Warren 2000; Machel 2004 for an up-to-date review) and all burial models, such as compaction flow, thermal convection, tectonic squeegee, topography driven flow, are essentially hydrological models, as they need to move large quantities of magnesium through the carbonate precursor. Massive dolostone bodies may have experienced more than one dolomitization episode which relate to different hydrologic regimes during their history. To reconstruct the diagenetic history and processes involved accurately, it is necessary to integrate petrographical and geochemical studies

with the sedimentary and structural evolution of the basin. This paper focuses particularly on the late burial precipitation of dolomite cements during collisional tectonics, and their relationship with the palaeohydrological regime developed in thrust-fold belt– foreland basin systems. Two hydrocarbon-bearing dolomite units involved in two remarkably different collisional systems are considered, that is, the Lower Jurassic carbonate platform of the central Southern Alps, northern Italy, and the Jurassic– Cretaceous slope deposits of central Apennines in central Italy (Fig. 1). The diagenetic study of the dolomitized bodies of the two areas, carried out on reservoir well cores and outcrop samples, indicates a multiphase dolomitization. A first phase of replacement occurred during the passive margin development, and was followed by the precipitation of different dolomite cements throughout the late stage of diagenesis related to continental convergence (Ronchi et al. 2003a, b, 2005; Murgia et al. 2004). We have paid particular attention to the dolomite

From: VAN BUCHEM , F. S. P., GERDES , K. D. & ESTEBAN , M. (eds) Mesozoic and Cenozoic Carbonate Systems of the Mediterranean and the Middle East: Stratigraphic and Diagenetic Reference Models. Geological Society, London, Special Publications, 329, 397–413. DOI: 10.1144/SP329.15 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Location map of study areas. (A) Central Southern Alps (B) central Apennines.

cement precipitation that constitutes only a few percent in volume of the entire dolostone bodies. These cements are mainly confined to fractured areas or dolostones with porous texture. These cements play an important role in the petrophysical properties of the dolostone reservoirs. In spite of similar petrography and tectonic setting, microthermometric data on fluid inclusions shows significant salinity differences of the diagenetic fluids circulating in the rocks during the collisional evolution of the two systems. Together with stable isotope analyses and burial modelling, our observations have led to an improved insight into the dolomitization processes in these systems.

Geological setting Southern Alps The Southern Alps tectonostratigraphic architecture is the result of a polyphase tectonics characterized

by Mesozoic east –west extension (Piedmont Ocean opening) followed by Cenozoic north – south compression (Adria and Eurasia plates convergence; Scho¨nborn 1992; Schumacher et al. 1997). During the rifting phase, from Norian to Middle Jurassic, the area was dominated by carbonate deposition (Fig. 2a). The few hundreds of metres thick carbonate platform sediments (which make up the Albenza Formation; Jadoul & Galli 2008) were flooded in Early Jurassic times and the area became an open carbonate shelf (Sedrina Limestone; Jadoul et al. 1994). In response to regional east –west oriented extensional tectonics, north – south elongated deep basins developed in the Lombardy Basin and were subsequently filled by pelagic cherty carbonate (about 1500 m thick Medolo Group; Castellarin & Picotti 1990; Gaetani et al. 1998). North–south palaeohighs persisted and thick slope breccia wedges developed at their margins. From Late Jurassic to Early Cretaceous

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Fig. 2. Tectonochronostratigraphic chart of the studied successions illustrating the differences in geodynamic setting through time. Tectonic phases shown on the right for: (a) Southern Alps; (b) central Apennines. The light grey fields with black rhombs refer to the stratigraphic position of dolomitized bodies. The dark grey fields represent evaporites.

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the basin underwent a post-rift thermal subsidence related to the drifting phase (Bertotti et al. 1993; Sarti et al. 1993; Ceriani et al. 2006 and references therein), and pelagic carbonate sediments accumulated in both structural basins and change in drowned highs. During Late Cretaceous, due to the onset of the north– south convergence phase and beginning of chain formation, a retro-wedge foredeep developed in the zone now belonging to the thrust-fold belts and was filled by thick flysch successions (Fig. 3a). During the Cenozoic, the southward migration of the Southern Alps chain caused a progressive southwards shift of the foredeep, pushing it to its present location in the Po Valley area; contemporaneously, the Cretaceous flysch depocentre area was first incorporated by advancing thrust sheets and then exhumed. The Po Valley is generally interpreted as a deformed foreland basin that suffered only moderate deformation in recent times (Miocene; Fantoni et al. 2004). In the context of this paper, dolomitization occurred in the Albenza Formation palaeohighs

and locally in the Sedrina Limestone and Medolo slope breccias adjacent to the palaeohighs (Fig. 3a).

Central Apennines As in the Southern Alps, the sedimentary successions of the central Apennines were deposited in a passive margin setting (Late Triassic to Cretaceous), that evolved to a pre-collisional convergent phase (Cretaceous to Oligocene) and was then involved into a collisional orogenic belt (Miocene; Marchegiani et al. 1999). Throughout the Triassic and lower Early Jurassic, shallow-water carbonates were deposited across the whole area. The Triassic peritidal dolomite platform has evidence of sabkha intervals and anoxic intra-platform basins, and an extensive carbonate platform (Calcare Massiccio Formation), similar to that of the Southern Alps (Albenza Formation) developed during the lower Early Jurassic (Ciarapica & Passeri 1998). During the middle Early Jurassic, platform stretching to the north caused the development of the Umbria-Marche

Fig. 3. (a) Lombardy Basin area: Southern Alps structural map with illustration of migration through time of the thrust front, distribution of dolomitization with respect to the Jurassic basins and palaeogeographic highs, and the location of studied outcrops and wells; (b) central Apennines structural map with indication of the Mesozoic– Cenozoic palaeogeographic units, the distribution of dolomitized bodies and the location of the studied outcrops and wells.

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Basin (UMB) with marly deposits and a long-lived carbonate platform to the south [Apulian and Apennine Platform realm (AAP); Santantonio 1993]. In between, the AAP –UMB slope transition area was dominated by resedimented facies composed of breccias and calcarenites. This palaeogeography lasted until the early Miocene when, in the Monte Vettore–Montagna dei Fiori area, the Apennine tectonics caused the formation of a NNW–SSE foredeep filled by turbidite deposits (Fig. 3b). During the early Pliocene, the AAP– UMB system experienced an orogenic peak and the zone was thrust and exhumed as the foredeep migrated eastward. The tectonic compression wave migrated northeastward until middle Pleistocene time, progressively involving the outer zone of the present-day Apennine chain (Centamore et al. 1992). In the central Apennine system, dolomitization occurs in the Lower Jurassic Calcare Massiccio platform of both the thrust-fold belt units and the foreland basin and in the Upper Jurassic slope facies of the Apulian Platform in the foreland succession.

Sampling strategy and methodology In order to investigate late stage dolomite precipitation in the two selected case studies, samples from both outcrops of the thrust-fold belts and subsurface of the related foreland basins were studied. In the central Southern Alps, samples from the Zandobbio area and from the Malossa and Seregna reservoirs were selected to represent the thrust-fold belt and foreland basin respectively (Fig. 3a). The Malossa field can be considered representative of foreland basin at the time of dolomite precipitation. Later it was involved in the frontal part of the thrustfold belt and, it now forms part of the deformed Southern Alps foreland basin. In the central Apennines, samples from the Montagna dei Fiori and Monte Vettore areas were studied in the thrust-fold belt, and samples from several wells distributed in the foreland area were analysed (Fig. 3b). The processes involved in the diagenetic evolution of the studied carbonate units were investigated through an approach that integrates petrographic and geochemical methods with thermal modelling. To define the paragenesis of studied rocks and the timing of precipitation of dolomite cements, a total of approximately 150 samples were observed both through transmitted light optical microscopy and cathodoluminescence (CL, using a Technosyn, model 8200MK2; 15 kV and 300–400 mA gun current); a few representative samples were observed also under scanning electron microscope (SEM), model Leika Cambridge Stereoscan 3608 with backscattered electron (BSE) detector.

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Carbon and oxygen stable isotope signature and fluid inclusions trapped in dolomite cements were studied to constrain the temperature and composition of dolomitizing fluids. Fifty samples from the Apennines and 74 from the Alps were analysed for stable isotopes. The analyses were generally performed on bulk samples but, where possible, microsampling of specific dolomite cements was carried out. Standard methodology was applied to the carbon and oxygen stable isotopes analyses: carbonate (4–6 mg) was allowed to react with concentrated orthophosphoric acid at 60 8C for dolomite samples. The CO2 content was measured in a Finnigan MAT 252 mass spectrometer and values are expressed as Peedee Belemnite formation international standard (PDB); results were expressed in parts per thousand (delta units). With this procedure no further correction for dolomite is needed. The petrography of fluid inclusions was studied on selected, unheated, and unstained thick sections (for a total of 26 samples), using an Olympus BX-60 microscope equipped with ultraviolet (UV) epi-illumination. Incident light fluorescence was used to distinguish petroleum fluid inclusions from aqueous inclusions. The fluorescence system used a 100 watt mercury bulb correctly focused and centred in its lamp housing, and using a narrowband ultraviolet (365 + 5 nm) excitation filter. Selected fields of view were stored digitally and printed as maps of fluid inclusion distributions. The microthermometric measurements were performed using a Linkam THMSG 600 heatingfreezing stage, calibrated with synthetic pure water and CO2 inclusions. Over 350 homogenization temperature (Th) and 380 final ice-melting temperatures (Tmice) were measured by cycling (Goldstein & Reynolds 1994). Special emphasis was placed on identifying fluid inclusion assemblages (FIAs), that is, the most finely discriminated groups of petrographically associated fluid inclusions, such as the inclusions from an individual growth zone or within a single healed microfracture which represent those inclusions trapped at about the same time (Goldstein & Reynolds 1994). Fluid-inclusion analyses were performed on dolomite cements representing the structural domains considered in both study areas (thrust-fold belt and foreland basin; Fig. 3a, b). All the data are taken only from primary fluid inclusions, that is, those inclusions clearly displaying a relationship to mineral growth. Figure 8a, b displays only data belonging to FIAs where it was possible to collect both Th and Tmice measurements on the same fluid inclusion. Afterwards, the thermal data were combined with the results of thermal modelling to define the age of dolomite cementation events with respect to the tectonic evolution of the studied areas.

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Paragenesis and dolomite precipitation In the samples from both studied areas, dolostones present the same variability in textures related to the limestone precursor. Fine-grained facies resulted generally in a massive replacive dolomite, represented by planar-s to non-planar texture (Sibley & Gregg 1987) with medium crystal size (from 50 –150 mm), whereas the grainstone/ packstone facies gave rise to sucrose dolomite with planar-e texture and coarse crystal size (150 – 500 mm). In the planar-s and planar-e textures a complex paragenetic evolution was reconstructed (Figs 4 & 5; Murgia et al. 2004; Ronchi et al. 2005). A multiphase dolomitization has been recognized consisting of a first replacement dolomite (D1), followed by several phases of dolomite cement precipitation (dolomites D2, D3, D4), interrupted by phases of dissolution and fracturing. † Dolomite 1 (D1) is represented by turbid dolomite crystal cores. In the Alps samples, the CL is orange whereas the Apennine samples vary from orange to violet, with a mottled aspect (Fig. 4b, e). Many small inclusions (few microns in size) and some traces of calcite have been detected by SEM analyses (Fig. 4f). Corrosion of the edges of the crystal cores are evidence of a dissolution phase, locally accompanied by fracturing. † Dolomite 2 (D2) is a cloudy dolomite cement with dull red/violet luminescence in CL. † Dolomite 3 (D3) corresponds to transparent dolomite cement overgrowth on Dolomite 2; it shows thin CL-zoning with a final bright orange or violet zone (Fig. 4). † Dolomite 4 (D4) includes saddle dolomite crystals (up to 600 mm) with curved crystal faces, cleavage and sweeping extinction (Radke & Mathis 1980); its CL appearance is dull dark red. D4 is present in pore spaces and fractures and constitutes the last carbonate phase in the Apennines’ samples. In the Alps, the saddle dolomite is rare and the last diagenetic phase consists of calcite cements (C1). The dolomitization was accompanied by a general increase in porosity with the formation of intercrystalline pores and enlargement of interparticle and vuggy porosity. This increase in porosity is attributed to the first replacement phase and subsequent dissolution. Dolomite cements partially fill the intercrystalline porosity and reduce fracture permeability that was formed partly during compaction and mainly related to tectonic deformation (‘over-dolomitization’; Lucia 2004). The massive dolomitization replacement phase (D1) appears to have been caused in a similar

way in the Southern Alps (Ronchi et al. 2005) as in the Apennines (Murgia et al. 2004), and its interpretation is based on its distribution. In the Southern Alps (Fig. 3a), dolostones are preferentially present in Lower Jurassic platform carbonates (Albenza Formation), that evolved in palaeohigh setting during subsequent extensional tectonism at the platform margin facing the lower Jurassic basins, and in breccia wedges at the platform/basin transition. In the Apennines (Fig. 3b), the Calcare Massiccio (Early Jurassic) is preferentially dolomitized in zones that evolved on palaeohighs in the Umbria Marche Basin; moreover, the dolomitization affected the calcarenite and breccia bodies lying on the flanks of the Jurassic/ Cretaceous basins. The fine-grained texture of the mud supported dolomitized facies and their relatively high oxygen isotopic values (around 0‰ PDB; Ronchi et al. 2005) suggest a low dolomitization temperature (,50–60 8C). Unfortunately, fluid inclusion analyses on D1 phases was not possible because of the scarcity of suitable inclusions in the mottled crystals: no microthermometry data could be collected. Based on these analyses, the replacement dolomites have been explained using the ‘compaction model’ (Fig. 6a, b; Machel 2004). In both areas the Jurassic/Cretaceous basin sediments could have provided dolomitizing fluids during burial compaction; fluid flow was funnelled in the more permeable bodies in basin flanks and palaeohighs and therefore, the source of Mg was most likely to have been sea water originally trapped in the pelagic sediments. In the Apennines, the Triassic carbonate-evaporate formations (Fig. 6b) could have expelled saline waters through synsedimentary extensional faults.

Stable isotopes Southern Alps The results of the analyses of the bulk samples from the central Southern Alps are shown in Figure 7. They are grouped with respect to the tectonic setting: 1) deformed foreland basin samples show a relative narrow range of carbon isotopic values (from þ2.5 to þ4‰ PDB), whereas the oxygen data are more scattered, ranging from 25.7 to 20.5‰ PDB; 2) the thrust-fold belt bulk samples are characterized by wide ranges in carbon (from þ1.5 to þ5‰ PDB) and oxygen (from 27.4 to 23.2‰ PDB) values. The isotopic values of the dolomite cements collected by microdrilling in the thrust-fold belt samples are included in a small field with d13C between þ3.5 and þ4‰ PDB and d18O from 27.7‰ to 28.2‰ PDB.

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Fig. 4. Photomicrographs of dolomitized facies. (a– c) Photomicrographs from Southern Alps samples: (a) planar light, coarse planar-s dolomite with intercrystalline porosity: the D1 is dark dolomite followed by cloudy (D2), transparent dolomite cements (D3), and saddle dolomite (D4). (b) the area of picture A under CL: note that the dark turbid rhombohedra nuclei (D1) are mottled and display a dark orange luminescence, the turbid dolomite overgrowth (D2) is dull and dark, D3 shows a thin bright outer zone, D4 is dark red. (c) SEM– BSE photomicrographs of D3 compositional zoning (Mn and Mg differences detected by XRF), and D2 with mottled aspect due to the presence of small calcite patches and inclusions. (d –f) Photomicrographs from central Apennines samples: (d) planar light, coarse planar-e dolomite with intercrystalline porosity; the rhombs nuclei and D2 cement are dark, while D3 dolomite cement is transparent. (e) detail of picture D under CL; D1 and D2 show red/violet luminescence, D3 is zoned bright violet. (f) SEM–BSE photomicrographs with D1 revealing calcite inclusions (white dots), D2 is homogeneously grey, in D3 very thin zoning is detected.

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Fig. 5. Paragenesis and relative timing of the main diagenetic phases common to the two studied areas. Early diagenetic processes and lithification are thought to have occurred in the marine to shallow burial environment; compaction, pressure-solution and fracturing are simplified; the dolomite replacement phase is interpreted as a burial process (that is, compaction model), followed by three phases of dolomite cement precipitation alternating with dissolution and fracturing events. The last cement phase is represented by a calcite spar filling voids and fractures in the Southern Alps domain, and by saddle dolomite (D4) in the central Apennines samples.

Central Apennines

Southern Alps

Stable isotope data for samples from the central Apennines are reported in Figure 7. The overall distribution of values for the d13C varies from þ1.5 to þ4.8‰ PDB in the whole set of samples. The dataset concerning the d18O can be differentiated according to the structural position of the samples (Fig. 3b): 1) the foreland basin samples are characterized by the highest d18O values (from 21.2 to 1.3‰ PDB); in the deformed foreland basin samples the range is intermediate (from 23.9 to þ1.5‰ PDB). 2) Samples from the thrust-fold belt show an oxygen ratio that is the most depleted with values ranging from 23.5 to 20.2‰ PDB.

The results of the analyses of fluid inclusions from the Southern Alps are shown in Figure 8a. They are collected with respect to the tectonic setting. (1)

Fluid-inclusion petrography and microthermometry Fluid inclusions found on the dolomite phases are mainly aqueous in composition, but petroleum fluorescent fluid inclusions were also observed. These hydrocarbon inclusions are both of primary and secondary origin. In this work, we have focused on primary aqueous inclusions, that is, those useful for diagenetic reconstruction. For each set of samples it was not always possible to collect data on all the crystalline phases composing the paragenesis. For this reason, diagrams in Figure 8 can report only data collected on a limited number of phases.

(2)

Thrust-fold belt samples contain primary fluid inclusions in dolomite phase D2 showing homogenization temperatures ranging from about 75 –100 8C with values more commonly around 80 –90 8C (mean Th 87 8C). Fluid salinity ranges from brackish to sea water (mean salinity 1.5 weight % NaCl equivalent). Calcite cements (C1) are also present. Nevertheless the scarcity of measurable fluid inclusions, homogenization temperatures (Th) range from 98–121 8C (mean Th 110 8C), and salinity close to that of sea water (mean salinity 3.2 wt% NaCl eq.). In foreland basin samples, D2 displays a scattered pattern in every fluid inclusion assemblage. This is due to re-equilibration of fluid inclusions as a consequence of thermal stretching and leakage-refilling. This process is likely to have occurred during flux of hot, low-saline fluids, probably those responsible for calcite (C1) precipitation. Despite the scattered distribution, the wedge-shaped dispersion of Th and salinity data converge on a point of low Th and low salinity roughly representing the original condition of the

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Fig. 6. Interpretative scheme for the pervasive dolomitization replacement phase (D1). (a) In the Southern Alps, the massive dolomitization is interpreted as being linked to the migration of fluids released during compaction of muddy carbonate lithologies in the Jurassic basins that migrate towards the palaeohighs: dolomitization was focused in the palaeohigh platform and along the porous slope breccia. (b) In the central Apennines the dolomitization in the palaeohigh and slope areas is linked to fluids released during compaction from Jurassic–Early Cretaceous carbonate pelagic basin sediments; some fluids may also have circulated along faults which accessed fluids from deeper formations.

dolomitizing fluid, with temperatures around 80 8C, and salinity between 1–2 wt% NaCl equivalent. Calcite (C1) in these samples displays similar high Th (mean Th 115 8C), as recorded in thrust-fold belt samples, and even lower salinity values (mean salinity 1.1 wt% NaCl eq.).

Central Apennines In the central Apennines, fluid inclusion data are reported in Figure 8b. (1) In thrust-fold belt samples, D2 Th values display large variations, ranging from 70 – 90 8C, with salinity ranging from 8–15 wt%

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Fig. 7. Stable isotope data of bulk dolostones and dolomite cements (stars), the latter being sampled by microdrilling.

NaCl equivalent. The wide salinity range can probably be explained through leakage and refilling of inclusions during the flux of later higher salinity fluids, possibly the same fluids responsible for precipitation of the later phases of cements (for example D4).

D4 dolomite shows the highest Th values (average value 119 8C) of the dolomite paragenetic sequence. The salinity for this phase ranges between 15.5–20 wt% NaCl equivalent, giving evidence of the entry of hot saline fluids.

Fig. 8. Fluid inclusion binary diagrams relating homogenization temperatures and salinity data for: (a) Southern Alps: 1) thrust-fold belt, and 2) foreland basin samples; (b) central Apennines: 1) thrust-fold belt, and 2) foreland basin samples. Diagram displays only data belonging to FIAs where was possible to collect on the same fluid inclusion both Th and Tmice measurements. Data are reported according to the mineral phase where they were collected; this was possible because every FIA in the same mineral phase displays the same trends. Star-point in Figure 8a 2 represents the minimum entrapment conditions for fluid inclusions entrapment in D2. Note different salinity of diagenetic fluids particularly evident in the thrust-fold belt samples.

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(2)

In foreland basin samples, D2 shows a spread in Th values, ranging from 75 –95 8C, with salinity ranging from 6–9 wt% NaCl equivalent. D3 precipitated from high temperature (Th from 95 –110 8C) and high salinity (between 6 and 11 wt% NaCl eq.) fluids.

Burial and thermal histories The burial and thermal history diagrams have been performed using the Eni proprietary Software GETw (1-Dimensional Integrated Basin Modelling System for Unix). The input variables required for this simulation contain many uncertainties that can influence the output results. These include the ancient heat flow, the surface palaeotemperature and palaeobathymetry, the lithology that controls compaction and thermal conductivity, the thickness and age of each sedimentary event, comprising especially the eroded sequence, etc. In the Southern Alps these uncertainties were limited by calibrating the thermal history with available data on organic matter maturity.

Southern Alps In the Southern Alps, the thermal history was calibrated using organic matter maturity data (Calabro` et al. 2003; Ronchi et al. 2003b; Fantoni & Scotti 2003; Scotti 2005; Bersezio et al. 2005; Ceriani et al. 2006; Zattin et al. 2006). In the two simulation locations, Malossa well and Zandobbio outcrop successions (Fig. 3a), we used the Mesozoic thermal regime evolution (heat flow applied at the base of the sediments) assessed by the above mentioned studies, based on the assumption of a heat flow rising during the Late Triassic –Early Jurassic time interval, due to crustal stretching (Fig. 9a3). The peak values of the heat flow were calibrated to the Bajocian/ Bathonian in age in several Mesozoic successions of the Southern Alps and seem to be reasonably uniform (ranging from 85 –105 mW m22; Fantoni & Scotti 2003). In this case study, a maximum heat flow of 95 mW m22 was used, based on the control points on the reliable maturity gradient measurements (vitrinite reflectance) collected in the nearby Iseo basin succession of Triassic to Cretaceous age (Calabro` et al. 2003; Ronchi et al. 2003b; Fantoni & Scotti 2003). Since the onset of the drifting phase (Bajocian/Bathonian) and following thermal subsidence, the heat flow progressively decreased, reaching the present day value of about 50 mW m22 (Fig. 9a3), as confirmed by the most recent reconstructions (Scotti 2005). In the foreland basin simulation point (Malossa), the maturity profiles suggest

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a different evolution of the thermal regime in recent times. In particular, a decreasing heat flow is assumed during the Oligocene– Miocene (Fig. 9a3) at the time of the formation of the strongly subsiding foreland basin. The heat flow decrease may be partially apparent due to a possible circulation of cool freshwater (coming from the Alpine orogen) into the deep carbonate series; the described hydrodynamic phenomenon likely took place in the Upper Miocene allowing thermal re-equilibration (Novelli et al. 1987). This supposed freshwater input is confirmed by the present-day occurrence of diluted waters in the Malossa reservoir, and evidence described in this paper. In the thrust-fold belt simulation location (Zandobbio, Fig. 9a1), the Early Jurassic succession was not deeply buried (and thus not significantly heated despite the high heat flow during the Jurassic) until Cretaceous times when it was buried under a thick sedimentary turbiditic succession, which was subsequently eroded. The thickness of Jurassic to Cenozoic sediments that produced the maximum burial and heating of the Early Jurassic carbonate is calculated from sedimentological evidence, and then was assessed by matching the resulting modelled thermal history with the observed organic matter maturity for the Early Jurassic sediments (measured Ro about 0.50 –0.55%). This corresponds to a temperature for the Albenza Formation of about 100 8C over several millions of years. The sedimentary sequence was involved in thrusting and exhumation with decreasing burial temperature, during the subsequent Alpine collision in the Cenozoic. The Foreland Basin simulation location (Fig. 9a2) is characterized by a moderate burial (1000 m/140 Ma) during Mesozoic –Eocene times, followed by a major burial (3500 m/40 Ma) during the Oligocene– Pleistocene times. The Dolomia a Albenza was heated progressively from 30 Ma reaching 140 8C today.

Central Apennines The thermal evolution of the study area is not well constrained by organic matter maturity data. The palaeogeodynamic position makes the occurrence of a very high heat flow in the Middle Jurassic less likely, mainly because of the greater distance from the rifting zone when compared to the Southern Alps (e.g. Stampfli & Borel 2004). A moderately high heat flow value (67 mW m22) has been assumed during the Bajocian –Bathonian for this location, followed by a decrease in heat flow to 46 –50 mW m22 that took place from the early Cenozoic (Fig. 9b3). In the thrust-fold belt area the heat flow remained at this value to the present

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Fig. 9. Heat flow and thermal history of (a) Southern Alps, and (b) central Apennines areas. In the Alpine area the heat flow evolution was characterized by high values during Jurassic time due to the Piedmont Ocean opening, while in the Apennines a minor heating is postulated. Due to the maturity data of organic matter a slight difference between thrust-fold belt and foreland basin heat flow is assumed. The thermal history of the dolomitized formations is the combination of heat flow evolution and burial history: highest temperatures are reached during periods of maximum burial under flysch deposits and therefore the timing of Tmax migrates laterally as the flysch belt migrates through time.

day; in the foreland basin location heat flow values decreased rapidly from the early Pliocene until the very low present day values (around 35 mW m22) were reached. The thrust-fold belt simulation location (Fig. 9b1) is characterized by a low burial (i.e. the top of Calcare Massiccio was never buried deeper than 2000 m) up to the early Miocene with consequent low heating. The top of this formation probably reached 2500 m of burial in the late Miocene resulting in burial temperatures higher than 80 8C. The foreland basin simulation location, where dolomitized slope sediments of Late Jurassic – Early Cretaceous age were studied, experienced a

similar burial and thermal history until the latest Miocene. Burial rapidly increased during the PlioPleistocene and carbonate formation temperatures increased up to 100– 110 8C (Fig. 9b2). The presentday heat flow used for the simulation has been derived from well data (BHT), and is caused by the thermal disequilibrium due to the high sedimentation rate: 2800 m in the last 4 Ma.

Fluid temperatures and timing of dolomite precipitation By matching the fluid inclusion precipitation temperature data with the thermal history of the

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host rocks the age of the dolomite cements precipitation can be reconstructed. The pressure correction was not applied here because the presence of methane gas saturation in the fluid inclusions is very likely and as a result the pressure corrections should be negligible and hence not applied in this discussion. The presence of hydrocarbons is supported by the occurrence of oil inclusions trapped in the dolomite cements and inferred from the geological context where the maturity of Triassic source rock has been reached in each petroleum system.

Southern Alps In the Southern Alps, the replacive dolomitization phase (D1) is interpreted as the result of a flux in the platform margin sediments of compaction fluids expelled from adjacent basinal sediments in a passive margin regime. This phase may have lasted for a few million years and was independent from the later thrusting. It can be tentatively put in a lapse time from the initial compaction and the time when the carbonate reached about 70 8C, minimum precipitation temperature for the D2 cement. According to the thermal histories shown in Figure 9a this interval covers the Late Jurassic – Late Cretaceous for the both thrust-fold belt and foreland basin areas. The precipitation of dolomite cements (D2– D3) likely occurred at between 80 –100 8C, as constrained by fluid inclusion microthermometry (Fig. 8a). These temperature conditions can be linked to the deep burial under the thick clastic sequence that accumulated in the South Alpine foredeep. In the structural domain that is now part of the thrust-fold belt this event took place during the Late Cretaceous. At that time, this domain was the retro-wedge foreland basin of the Alpine chain where thick Cretaceous flysch units were deposited; the Cretaceous flysch (Fig. 2a) depocentre extended along the present day east –west narrow belt from lake Iseo to Lake Como (Fig. 3a). In the thermal history diagram (Fig. 9a1), the 80 8C isotherm is crossed in the uppermost Cretaceous (70 Ma). Subsequently, burial and temperatures continued to increase to a maximum that was recorded by the calcite cements at temperatures around 110 8C (50–60 Ma). Finally, thrusting and exhumation of the area took place, causing cooling. In the deformed foreland basin the temperatures recorded by the D2 dolomite (average 85 8C) were reached later (i.e. around 25 Ma, Fig. 9a2), corresponding to the foredeep formation in front of the Miocene thrust belt (collision phase of Fig. 2a, south to the Miocene thrust front of Fig. 3a). As in the thrust-fold belt, even in the foreland basin samples, the calcite records the temperatures of maximum heating (Fig. 9a).

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Central Apennines The overall diagenetic evolution of the central Apennines is similar to that described for the Southern Alps. In the central Apennines, the replacive dolomitization phase (D1) occurred during an extensional tectonic regime in Late Jurassic–Early Cretaceous (Fig. 2b). Matching fluid inclusion Th data for the D2 dolomite phase with the thermal histories of Apennines thrust-fold belt and foreland basin (Fig. 9b), it is evident that this event occurred later than in the Southern Alps domain accordingly to the timing of the onset of collision in the two orogens (Fig. 2a, b). In the Apennines the D2 precipitation took place earlier (around 10 Ma) in the thrust-fold belt than in the foreland basin (around 2 Ma), in the Calcare Massiccio and earliest Cretaceous slope sediments respectively. In the foreland basin, D3 dolomite was precipitated more recently in the studied formations when the deepest burial conditions were reached. In the thrust-fold belt unit, D4 dolomite records a hydrothermal pulse with precipitation temperature higher than those reached during burial (Machel & Lonnee 2002).

Composition of diagenetic fluids The carbon isotopic data of the dolomite ranges between þ1.5 and þ5‰ PDB, that is, in the range of the Jurassic marine carbonate signature (Allan & Wiggins 1993). This suggests that dolomitization occurred in a diagenetic environment with a high rock/water ratio and therefore the signature was inherited from the original carbonate sequence. The oxygen isotope data are more scattered; however, the values from Southern Alps samples were always more depleted than the values obtained from the central Apennine samples. Moreover, in each area, a trend of progressively increasing values was observed from the thrust-fold belt toward the foreland. The most depleted values were observed from dolomite cements sampled using a microdrilling device. The oxygen isotopic depletion can be the result of multiple causes. Those considered most important are the temperature of crystallization and fluid isotopic composition. The more depleted values of the thrust-fold belt areas in comparison to the values obtained from the foreland basin samples may be due to higher temperatures and/or longer time of residence of the carbonate at high temperature. The significant difference between d18O from samples from the Southern Alps and central Apennines may be due to either temperature fractioning or to the composition of the diagenetic fluids (Land 1985). It was possible to infer the composition of the diagenetic fluids (Allan & Wiggins 1993) by

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Fig. 10. d18O of dolomite cements v. temperature using Th values determined through fluid inclusion microthermometric analyses. The correlation between the two set of values can be interpreted in terms of fluid isotopic composition according to the diagram proposed by Allan & Wiggins (1993). The Southern Alps samples suggest involvement of diagenetic waters compatible with sea water or diluted sea waters, while the Apennine samples may suggest formation brines.

integrating these isotopic data with the Th values observed from fluid inclusion analyses. This is shown in Figure 10. The dolomite cements d18O of Southern Alps samples (around 28‰ PDB) matched with the fluid inclusions Th (mean value

87 8C), give an oxygen isotopic nature of the fluid ranging between þ1 to 23‰ SMOW. This SMOW value indicates a low-salinity fluid, which is consistent with salinity measured in fluid inclusions (Land 1985). In the central Apennine samples the less depleted oxygen values (dolomite cements estimated around 24‰ PDB) linked to the fluid inclusions Th data indicate a SMOW ranging from þ2 to þ7‰ PDB. This range of values is characteristic of formation brines (Fig. 10). These data demonstrate a remarkable nature difference between the samples from the Southern Alps, which are characterized by low saline diagenetic fluids, and the central Apennine samples, which yielded values typical of saline brines. This difference was observed despite the reconstructed diagenetic evolutions of the areas being similar and the areas being situated in similar geodynamic contexts.

Dolomite cementation model

Fig. 11. Simplified scheme (not to scale) illustrating the different structural, morphological, tectonic and climatic settings in the Southern Alps (a) and central Apennines (b) thrust-fold belts during the studied diagenetic events. See text for explanation of the suggested controlling factors. These regional effects had an influence on the nature and extent of late dolomitizing diagenetic fluids in the areas studied.

The dataset reported is consistent with a paragenetic history where the dolomite cements in the Southern Alps and central Apennines precipitated under deep burial conditions, during collisional tectonics. In this context the driving flow mechanism is the tectonic loading which forced pressurized fluids to move from the collision zone to the foreland (Oliver 1986). A dolomitization process often referred to this environment is related to tectonic (squeegee) fluid expulsion (Machel 2004), and many high temperature dolomitized bodies,

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frequently associated with Pb –Zn deposits, have been interpreted through this process (Montanez 1994; Machel et al. 2000). The diagenetic fluids are usually saline brines consistent with mixing with metamorphic fluids; nevertheless topographydriven flow may import meteoric water into the system, thus providing a mix of different waters which results in a variety of final diagenetic fluids. The dolomite cements analysed during this study can be easily related to a consistent tectonic model of dolomitization, despite diagenetic fluids having had strikingly different compositions; that is, there are highly saline pore fluids in the central Apennines and pore fluids of low salinity in the Southern Alps. The occurrence of diluted diagenetic waters in the Southern Alps case is consistent with the present-day occurrence of brackish waters in Mesozoic carbonates in the foreland basin subsurface reservoirs, at more than 5 km depth. The dolomitization potential of these low saline fluids, however, remains to be demonstrated. In the Southern Alps samples, it is proposed that this can be explained by the magnesium enrichment of the fluid through rock-water interaction. This can be indicated by existence of dissolved edges of the dolomite rhombs of D1. The subsequent D2 cements precipitation could have happened during the rise of temperatures which favoured precipitation instead of dissolution.

Discussion and conclusions Despite the similar tectonostratigraphic evolution of the dolomitization fluids involved in the precipitation of dolomite cements in the two studied areas, the resultant dolomite cements have been shown to be different in composition. In the central Apennine samples, an increase in salinity in the cement paragenesis has been noted, whereas in the Southern Alpine samples a decrease in salinities from sea water to brackish fluids in the cements have been recorded. This opposite trends, that is, the difference of fluids responsible for late dolomite precipitation, is interpreted in terms of diagenesis during different phases of thrust-belt morphology–tectonic evolution. This interpretation is supported by several lines of evidence: † The Alpine chain was widely emerging since the latest Cretaceous, as testified by the huge amount of detrital materials produced by its erosion (Cretaceous flysch). Therefore, the occurrence of a large recharge area for meteoric waters circulating in the thrust-belt foreland system is likely. † During Cenozoic, central Alps should have developed a rapidly rising, high relief collisional

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belt which acted as the source region for foreland basin deposits. This fact is consistently supported by the large volume of detrital material deposited on the foreland basin (Kuhlemann et al. 2002; Spiegel et al. 2004) and the coarse texture and mature composition of the detritus (Carrapa & Di Giulio 2001). † The high relief of the Alpine belt, probably caused a topographic barrier for atmospheric circulation increasing rainfall and, in turn, meteoric water supply to the hydrogeological system developed in the belt. This may have caused a high hydraulic head which allowed surface water to penetrate deep into the Mesozoic carbonates of the coeval foreland basin. † The central Apennines area is thought to have been barely emergent during Miocene –Pliocene time as inferred from the very few erosional products that it sourced. The turbidites that accumulated in the foreland basin were fed longitudinally from the NW and the sedimentary contribution to the basin fill of the central Apennines belt is considered to be negligible (Chiocchini & Cipriani 1989). Therefore, the recharge area for meteoric waters circulating in the thrust-belt foreland system can be expected to be limited. In addition, the low topographic relief of the belt probably was insufficient to create a topographic barrier to atmospheric circulation. † If this were the case, then the rainfall in the area would be limited. The low relief could have also caused a low hydraulic head which would not have allowed surface water to penetrate deeply into the Mesozoic carbonates in the foreland basin. As a result, high salinity waters fed by deeper evaporite-bearing stratigraphic levels (such as the Burano Formation) which also acted as the de´collement level during thrusting, are considered to have been the dominant source for diagenetic fluids in the central Apennines. The same squeegee mechanism valid for the Alps may be interpreted as the driving mechanism here for fluid flow. The two study cases indicate that collision tectonics is a driving process for late dolomitizing events in carbonate units involved in thrust-fold belt – foreland basin systems. These results also show that the diagenetic pathways followed by dolomitized rocks in the different part of the system were remarkably similar, despite the fact that diagenesis occurred at different times following the onset of the tectonic deformation. In other terms, it testifies of a sort of ‘diagenetic wave’, coupled to a ‘tectonic wave’, migrating in accordance to the tectonic evolution of the collisional systems.

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In conclusion, this comparison between the two study areas highlights the importance of the subaerial relief of the collision belt in controlling large-scale hydrogeological fluid circulation of the system. This in turn controls the composition of dolomitizing fluids circulating at different depths. The importance of relative relief in collision belts in the control of the hydrodynamic of the thrust-belt and foreland basin is well established. However, the potential control, on the composition and distribution of dolomitizing fluids within these tectonic settings has been previously overlooked in the construction of diagenetic models to explain regional dolomite precipitation. The research is part of an Eni R&D project dealing with diagenesis in carbonate reservoirs. We thank all the Eni colleagues that participated to the project and particularly T. Ricchiuto for the stable isotopes data, R. Fantoni and R. Longoni for structural geology and basin analysis support. We also acknowledge F. Jadoul for fruitful discussions on stratigraphy and the geological model. Financial support was provided by Eni and FAR (Pavia University). We also acknowledge K. D. Gerdes, R. Swennen and P. Wagner for their contributions that improved the original manuscript.

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Index Page numbers in italic denote figures. Page numbers in bold denote tables. Abu Dhabi platform 133, 135– 137 accommodation space regional control 11, 14– 15 Sarvak Formation 209– 211 TSPs 17, 20–36 Adam Foothills outcrops 164, 168 Adayah Formation 48, 51 Adnet reef 293, 294, 313 African Plate 10 Afro-Arabian Plate 27– 31 Neotethys border 18, 20– 23 age dating, Dezful Embayment 252–253 Agoudim Formation 67, 68, 76– 77 Ahmadi Member 359, 360, 365 Ahwaz Sandstone Member 223, 238, 239, 240 Akhdar Group 293, 294 Al Daww depression 44, 45, 61 Albian Bangestan Group 255–263 Basque-Cantabrian platform 317– 340 algal facies, Miocene, Ermenek Platform 275, 276, 277, 278, 279, 280, 282, 283, 285– 286 algal mounds, Cretaceous, Oman 158 Aliaga outcrop 116, 117, 123 sequence stratigraphy 128, 129– 131, 134 Aliaga-Miravete anticline 114, 116, 118 Allan/Muss Formation 48, 51, 58, 59 Alps, Southern burial and thermal history 407 diagenetic fluids 409– 410 dolomite cement precipitation 402, 403 modelling 410 –411 timing 409 fluid inclusions 404–405, 406 geology 398– 400 stable isotopes 402, 409 stratigraphy 399 Amanus Shale Formation 46, 48, 50, 51, 55 biostratigraphy 49 hydrocarbon systems 60–61 lithology 49 nomenclature 47 Amellago Formation 67, 68, 76–77 Amellago ramp system 67–85 biostratigraphy 72, 73, 76, 77 comparison with Persian Gulf 81, 84 comparison with West Caicos 84 cycles 76–77 facies distribution 78–81, 82, 83 evolution, large scale 76– 77 hemipelagic basin facies 69, 73, 76 inner ramp facies 69, 71, 74, 82 middle ramp facies 69, 72, 74– 75, 82 ooid-free muddy ramp facies 69, 76, 82 outer ramp facies 69, 73, 75–76, 82 Amellago transect 66, 67, 68, 76 Ammonitico Rosso 22, 25

Anaran Anticline 344, 345, 346 dolomitization 347, 348 faults and joints 349, 351– 353 regional comparison 388–392 reservoir modelling 382, 384– 387 stratigraphy 353 –363 structure 349– 353 uplift 382 anhydrite Asmari Formation 233– 236, 237, 240, 250, 251, 252, 257 Gachsaran Formation 236, 240, 252 see also Butmah Formation; Kurrachine Anhydrite Formation; Sergelu Formation anoxia see oceanic anoxic events (OAEs) Apennines, Central 398 burial and thermal history 407– 408 diagenetic fluids 409–410 dolomite cement precipitation 402, 403 modelling 410– 411 timing 409 fluid inclusions 405– 407 geology 400–401 stable isotopes 404, 409 stratigraphy 399 Aptian Khami Group 353–355 platform-basin transition, Galve sub-basin 113 –139 Aptian, Early, OAE 27, 132, 135 Aptian, Late unconformity, TSPs 11, 13, 26, 28, 25– 29 Aptian, Middle, transgression 25, 26 Apulian carbonate platform 31 Aquitanian, Asmari Formation 251–257 Arabian Plate 10 palaeogeography 222 tectonic setting 43, 45–46, 189 Triassic deposits 46–47 Arabian Platform, TSPs 18– 36 aragonite dissolution 379 –380 Artoles Formation 116, 118 chemostratigraphy 123 sequence stratigraphy 125 Asmari Formation 219, 220, 222, 223 anhydrites 233– 236 biostratigraphy 240, 243– 246 clinoforms 231 lithostratigraphy 236, 237, 238, 240 sedimentology 226– 236 Sr isotope stratigraphy 240– 245 Asmari Reservoirs 219, 257, 258, 259, 260 Atlantic Ocean, rifting 18–27 Austria, Adnet reef 293, 294 Austrian Event 29 Bab Basin 149, 150 Bab Member 149, 150 Bajocian, TST 29

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Bangestan Group 190, 192, 193, 207, 347 dolomitization 344 stratigraphy 353, 354, 355–363 structure 351 Barremian, Upper, carbonate platforms 118, 135, 138 Basal Anhydrite, Asmari Formation 223, 233–236, 237, 250, 251, 256, 257 basin, intra-shelf, Sarvak Formation 212 basins, accommodation space 11, 14–15 Basque-Cantabrian Basin 317–340, 318, 319 bathymetry, and accommodation space 11, 14– 15 Benasal Formation 116, 118 biostratigraphy 123 carbon isotope values 125 depositional model 120 sequence stratigraphy 130– 131 Berriasian, Lower, Iberian Basin, epeiric carbonate ramp 93, 94, 101, 108 Bibi Seydan outcrop 190, 194, 197– 200, 212 biostratigraphy Aptian Galve sub-basin 120–123, 137–139 Khami Group 353 Carnian, Palmyrides megacycle B 53, 56, 58 Jurassic, oolitic ramp, High Atlas 72, 73, 76 Miocene, Ermenek Platform 270, 275 –286 Norian, Palmyrides megacycle C 58– 60 Oligo-Miocene, Asmari Formation 227, 228, 240, 243–246, 249, 251, 252–253 Rhaetian, Palmyrides megacycle D 60 Scythian, Palmyrides megacycle A 49, 53 bioturbation, Sarvak Formation 361 biozones, Oligo-Miocene, Dezful Embayment 244, 245– 246, 252– 253 bitumen, Gadvan Formation 375 Bonarelli Event 27 brecciation Anaran Anticline 369, 370, 371–372, 372, 375, 377 Mahil Formation 298–299, 300, 301, 305 Breistroffer Event 159 Bundsandstein Group 293–297 stable isotope data 307, 308–309 Burdigalian Asmari Formation 252, 253, 255, 256, 257 Ermenek Platform 268, 269, 270–271, 286 burrow fills, dolomitized 363– 365 Butmah Formation 48, 51, 58, 59 calcite, blocky 296, 300, 301, 304 calcite compensation depth 33 caliche, Mahil Formation 298, 300, 301, 305 Calvo¨rde Formation 293, 294 meteoric diagenesis 294–297, 298, 311 stable isotope data 307, 308–309 Camarillas Formation 116, 118 Canada, fracture-controlled dolomite 390, 391–392 carbon isotopes Galve sub-basin 123, 131, 133, 134 Ramales platform 334–335 carbonate factories Cretaceous, Oman 158, 159 Jurassic, Amellago ramp system 81 Triassic 20 see also carbonate production

carbonate platforms Arabian Plate 46 Basque-Cantabrian Basin 317–340 epeiric 89, 145, 146 Oman 146– 160 channelized systems 164–184 Ermenek, Turkey 265–288 Kangan 18 Majorca 286 –287 Mut Basin 286–287 Sarvak Formation, Iran 187–214 see also platform-basin transition carbonate production Cenomanian– Turonian, Sarvak Formation 199– 200, 203– 204, 210– 213 Miocene, Ermenek Platform 285 –286 Oligo-Miocene, Dezful Embayment, fauna 227, 228 see also carbonate factories carbonate ramps 18, 23–26 epeiric, Iberian Basin 89–108 Kimmeridgian 93, 94, 96, 100–101 Oxfordian 93, 94, 95–100 Tithonian–Berriasian 93, 94, 96, 101, 103 oolitic Amellago system 67–85 High Atlas 65–85 carbonate sequence development 11, 14– 15 Sarvak Formation 209– 211 carbonate-siliciclastic platform, Aptian, Galve sub-basin 113– 139 Carnian, Palmyrides megacycles 53–60 Carnian Salinity Crisis 47 Castellar Formation 116, 118 cement Anaran Anticline, stratigraphy 379–382 blocky calcite 296, 328 calcite 326–327, 330, 379 carbonate 328, 331 dolomite, late stage precipitation 397– 398, 409 fibrous 294, 296, 306, 324, 325, 336 –337 Ramales platform 324– 331, 338 –339 saddle dolomite 371 –379 spar calcite 327–328, 330, 331 vadose sediment 379, 381 Cenomanian –Turonian, platform-basin transition, Sarvak Formation 187–214 Cenozoic, near-base unconformity, TSPs 11, 13, 30, 29– 33 Cenozoic tectonostratigraphy 13 Central Atlantic Magmatic Province LIP 20 chalk, Oman epeiric carbonate platform 158 channel incisions Natih Formation 156, 159, 165, 167, 169–171 differentiation from tidal channels 181 –183 incision fill 169, 171, 183 reservoir heterogeneity and prediction 183–184 channelized systems Natih Formation 164–184 see also channel incisions; channels, tidal channels, tidal Natih Formation 165, 167, 170, 171–181, 172–181 channel fill 181, 184 differentiation from incisions 181 –183

INDEX reservoir heterogeneity and prediction 183, 184 top of incision-fill 171– 172 Chaotic Unit 204, 205 Chattian, Asmari Formation 249 –250, 252 –253, 254, 255, 256 Chert Formation 116, 118 carbon isotope values 123, 125 depositional model 120 orbitolinids 122, 123 sequence stratigraphy 125 Chidan outcrop 225, 235– 236 coral buildups 252 Sr isotope stratigraphy 241, 242, 243 climate, early Aptian 113 climate change 11 cyclicity, ooid production 84–85 meteoric diagenesis 312 Oligo-Miocene, Dezful Embayment 254 clinoforms Amellago transect 80, 81 Asmari Formation, Tang-e-Gurguda outcrop 227, 231, 246, 250, 254 Sarvak Formation 198– 200, 203, 205–206 compaction, meteoric diagenesis 294, 302 Conchodon Formation 407 coral buildups Asmari Formation 227, 229 Miocene, Eremenek Platform 275, 276, 277, 278, 280, 282, 283, 286 coral reefs Aptian, Galve sub-basin 120, 121, 124, 137– 138 Kimmeridgian, Iberian Basin 106 see also Tropfbruch coral limestone Cretaceous channelized systems, Natih Formation 164–184 epeiric carbonate platform, Oman 146–160 near-base unconformity TSPs 11, 12, 24, 26, 23–27 Cretaceous, Lower Basque-Cantabrian platform 317– 340 Galve sub-basin 113 –139 Cretaceous Magnetic Quiet Zone 27 Cuadro marls 321 Cueto Misario section 327 Dachstein Platform 293, 294 Dariyan Formation dolomitization 369 saddle dolomite 371, 372 stratigraphy 354–355 tidal shoal 373, 374 Dashtak Formation 27 Deccan Traps LIP 31 Dezful Embayment 219– 260, 221 biostratigraphy 240– 246 revised zonation 244, 245– 246, 252– 253 controlling factors 253–254 facies substitution diagram 226 lithostratigraphy 223 sedimentology 226–236 sequence stratigraphy Miocene 250– 252 model 256–257

Oligocene 246– 250 previous interpretations 254–256 see also Asmari Formation diagenesis fluids, Alps and Apennines 409–410 hydrothermal stage 339 late burial 380, 382 meteoric control mechanisms 311–313 intermediate burial 338–339 karst unconformities 291–313 phreatic 337–338 shallow burial 338 Ramales platform 324–331, 335–340 seafloor marine 336–337 shallow burial 379–380 uplift-related 382 dissolution aragonite 379 –380 meteoric diagenesis 294, 296, 302, 305, 306 Ramales platform 324, 331, 333, 338– 339 dolomite cement precipitation Alps and Apennines 397–412 modelling 410–411 timing 409 matrix replacement 363– 371, 380, 402 Ramales platform 328, 330, 331, 339 saddle 371–378, 380, 381, 402 fracture-controlled Canada and USA 391–392 Gulf of Suez 389–391 reservoir modelling 382, 384–387 Zagros fold belt 388–389 dolomitization 397 fracture-controlled Anaran Anticline 343– 393 regional comparision 388–392 Mahil Formation 298– 299, 300, 302–304 El Garia Formation 35 El Perello sub-basin 115, 118 Ermenek Platform 265–288, 266, 268– 269 biostratigraphy and lithostratigraphy 270 comparision with other Miocene platforms 286 –287 controlling factors 287 –288 depositional models 280 facies associations 275, 276, 277–278, 281, 282 large scale evolution 285– 286 sequence stratigraphy 278– 283 stratigraphic architecture 269, 270–283 Langhian 271 –283 Escucha Formation 116, 118 eustacy 11 accommodation space 16–18, 23– 25, 29– 31 cyclicity, ooid production 84 see also sea level change evaporites Arabian Plate, Triassic 46–47 Messinian Salinity Crisis 35 Palmyrides megacycles 49, 50, 51, 54, 56, 60 TSTs 18, 28 extinction 11, 21, 28

417

418 faults, Anaran Anticline 349– 353, 366, 369, 370– 371, 372, 378 fauna see biostratigraphy folding, Anaran Anticline 349, 353, 371 foraminifera Albian, Ramales platform 322, 323, 324 Miocene, Ermenek Platform 275– 283, 285– 286 Forcall Formation 116, 118, 122, 125 depositional model 120 sequence stratigraphy 129, 130 fractures dolomitization, Anaran Anticline 343– 393 meteoric diagenesis 296 Mahil Formation 301, 305 Ramales platform 328, 333, 338 Tropfbruch coral limestone 302, 306, 310 Fuente Frı´a section 325, 327 Gachsaran Formation 233, 235, 236, 240, 252 Sr isotope stratigraphy 243 Gadvan Formation 353– 354 dolomitization 366 saddle dolomite 371, 372 hydrocarbons 375 Galve sub-basin Aptian platform-basin transition 113– 139 biostratigraphy 123, 134, 137–139 chemostratigraphy 123, 125, 131, 133, 134 facies analysis 118–123, 119 geological setting 114, 115, 116, 118 HSTs 129– 131 lithostratigraphy 118 LST 130 sea level change and deposition 133, 135 –137 sequence stratigraphy 123– 131, 136 TSTs 127, 129–131 Germany, Rogenstein 293 –297 Ghirab Member 370 Gialo Formation 34 greenhouse gases 11, 21 Gulf of Suez, fracture-controlled dolomite 389–391 Gurpi Formation 233 Hammam Faraun Fault, fracture-controlled dolomite 389–391 Hawar Member 148, 157 Hayan Block 43, 49, 54 heat flow, Alps and Apennines 407–408 Hercynian unconformity 18– 19 High Atlas mountains geology and tectonics 66–67 oolitic carbonate ramp 65–85 high stand systems tracts (HSTs) 15, 129 High Zagros 189, 190, 221 hydrocarbon reservoirs channelized systems 183–184 Oligo-Miocene, Dezful Embayment 219–260 Triassic, Palmyrides 43, 44, 60–61 TSTs 15 hydrocarbons, NW Dome 375 Iberian Basin, Upper Jurassic epeiric carbonate ramp 89–108 geology 90, 91, 92, 93

INDEX sedimentary evolution 93–103, 106–108, 107 sequence stratigraphy 92, 93–103, 94 Iberian Massif 90, 91 Ilam Formation 351, 360, 362–363 incisions see channel incisions inclined ramp, carbonate deposition model 16 inclusions, fluid Central Apennines 405–407 Southern Alps 404– 405, 406 intra-Oxfordian unconformity TSPs 11, 12, 22, 24, 20–25 Iran Dezful Embayment 219– 260 Sarvak Formation 187– 214 Zagros Mountains fracture-controlled dolomitization 343–393 geology 189– 190 iron sulphides, Anaran Anticline 371, 372 Izeh outcrop 225 Sr isotope stratigraphy 241, 242, 253 Izeh Zone 220, 221 Jabal Akhdar outcrop 164 incision-fill 169, 172 Jabal Madar outcrop 164, 168 bioclastic tidal channels 172–173, 174, 175 channel incisions 169, 170, 171 Jabal Madmar outcrop 164, 168 bioclastic tidal channels 175, 176, 177, 178 channel incisions 169, 170, 171 Jabal Nadah outcrop 164 Jabal Qusaibah outcrop 164, 168 bioclastic sandwave complex 177, 178, 179 bioclastic tidal channels 173 Jabal Salakh outcrop 164 Jabal Shams outcrop bioclastic sandwaves 179, 180, 181 ‘canyon’ 179, 180 channel incisions 169, 171 Jahadabad outcrop 225 Sr isotope stratigraphy 241, 242 Jahrum Formation 223, 239, 240, 246, 249 Jihar Fault 45 joints, Anaran Anticline 349–353, 368, 370– 371 Jura Platform 106– 107 Jurassic, earliest, unconformity, TSPs 11, 12, 19, 20, 21 Jurassic, Lower-Middle, carbonate ramp, High Atlas 65– 85 Jurassic, Upper, epeiric carbonate ramp, Iberian Basin 89– 108 Karoo-Ferrer LIP 22 karst Dariyan Formation 354– 355 Mauddad Member 356– 357 Sarvak Formation 361– 362 unconformities, meteoric diagenesis 291, 306 karstification, Tropfbruch coral limestone 302, 303, 306 Katoola outcrop 225, 232, 233 lithostratigraphy 236 Sr isotope stratigraphy 241, 242, 243

INDEX Kazerun Fault 190, 194 Kazhdumi Formation 29, 191 correlation with Natih Formation 207, 208, 209 sequence heirarchy 206 stratigraphy 192, 193, 207, 208, 354, 355 –357 Khami Group 347 dolomitization 344 stratigraphy 353–355 structure 351 Kharaib Formation 147, 148, 151, 157 Khuff Formation 20 Kimmeridgian, Iberian Basin, epeiric carbonate ramp 92, 93, 94, 96, 97, 100–101, 102, 106– 107 Kimmeridgian transgression 26, 25 Ko¨selerli Formation 269, 270 Kuh-e-Asmari outcrop 225, 233, 234, 235 lithostratigraphy 236 Sr isotope stratigraphy 241, 242, 243 Kuh-e-Bangestan outcrop 225, 233, 234 Sr isotope stratigraphy 241, 242, 243 Kuh-e-Khami outcrop 225 Sr isotope stratigraphy 241, 242, 243 Kuh-e-Khaviz outcrop, coral buildups 229, 250 Kuh-e-Landareh anticline 190, 194 Kuh-e-Pashmi dolomite plume 367, 368, 369, 370, 378 hydrocarbons 375 Kuh-e-Razi outcrop 225 coral buildups 229, 246, 250 Sr isotope stratigraphy 241, 242, 243 Kurrachine Anhydrite Formation 18, 46, 48, 49, 54, 55 nomenclature 47 Kurrachine Dolomite Formation 43, 48, 49, 50, 51, 55, 56, 57 biostratigraphy 53, 56 hydrocarbon systems 60–61 lithology 56 nomenclature 47 Langhian, Ermenek Platform 268, 269 stratigraphic architecture 271– 283 large igneous provinces (LIPs) 11, 20, 21, 25, 27, 31, 131 Las Parras sub-basin 115, 118 lead ore, Ramales platform 331 limestone, coral, Tropfbruch 293, 294, 301, 302, 303, 304, 305–306, 307 Lombardy Basin 400 Maastrichtian transgression 29, 30 Maestrat Basin 116, 130 Mahil Formation 293, 294 meteoric diagenesis 298– 305, 313 stable isotope data 305, 307, 309 Majorca, carbonate platform 286– 287 Marrat Formation 20 Mauddud Member 350, 356– 357, 358 karst dolomitization 368– 369 maximum flooding surfaces (MFSs) 46 megacycles, Triassic, Palmyrides 43, 48, 49–60 Mesozoic, tectonostratigraphy 12–13 Messianian Salinity Crisis 35

419

Messinian transgression, TST 11, 35, 36 unconformity, TSPs 11, 13, 34, 33–35, 36 micritization, Mahil Formation 301– 302, 303 microbialites Calvo¨rde Formation 293, 295, 297 Galve sub-basin 137– 138 Middle Anhydrite, Asmari Formation 223, 233, 235–236, 237, 250, 252, 257 Milankovitch cycles 311 mineralization, Ramales platform 331, 339 Miocene Ermenek Platform 265 –288, 266, 269 sequence stratigraphy, Asmari Formation 250– 252 transgression 36, 37 see also Oligo-Miocene Miravete anticline 114, 116, 118 Miravete fault 118 Miravete outcrop 116, 117, 123, 124, 134 chemostratigraphy 123 correlation 132 sequence stratigraphy 126, 129 –131 monsoon systems, meteoric diagenesis 312 Morella Formation 116, 118 Morella sub-basin 115, 117 Mut Basin 266, 267, 270, 286–287 Mut Formation 269, 270 Muti Formation 169, 172 Nahr Umr Formation 148, 152, 159 Natih Formation 146, 189 channelized systems 164–184 correlation with Sarvak-Kazdumi sequence 207, 208, 209, 212 sequence stratigraphy 165– 169, 208 Natih-e Member 146, 147, 148, 150, 153, 154, 155, 156 channel incisions 156, 159 correlation 153, 154, 157, 209 late Albian OAE 159 palaeogeography 153 sea level change and deposition 157–160 seismic data 155, 156–157 Natih-f Member 157 Natih-g Member 157 neomorphism Ramales platform 324, 337 Tropfbruch coral limestone 305– 306 NeoTethys Ocean 19–20, 46 margins 18, 189 Norian, Palmyrides megacycles 58– 60 NW Dome 348, 351, 354 dolomitization 369, 370, 372, 375, 378 hydrocarbons 375 occlusion, pore systems 294, 295–296, 302, 305–306 oceanic anoxic events (OAEs) 11 Early Albian 27 Early Aptian 25, 131, 132, 159 Triassic– Jurassic boundary 20 Ojebar section 321, 322, 323, 327 Oliete sub-basin 115, 116 Oligo-Miocene, Dezful Embayment 219–260

420

INDEX

Oligocene sequence stratigraphy, Asmari Formation 246– 250 unconformity, TSPs 11, 13, 32, 34, 31– 34 Oman, epeiric carbonate platform 146–160, 147 geological setting 146–148 sea level change and deposition 157–160 seismic data 150– 157 Oman, Mahil Formation 293, 294, 298–305 stable isotope data 307, 308 Oman, Natih Formation 146, 189 channelized systems 164–184 Oman platform, sea level change and deposition 133, 135– 137 oncoids, Tithonian, Iberian Basin 107–108 Ontong Java LIP 27, 131 ooids Amellago ramp system 65– 85 climate-driven cyclicity 84–85 dissolution, Calvo¨rde Formation 294, 298 eustacy-driven cyclicity 84 oolite 29 Calvo¨rde Formation, meteoric diagenesis 293–297, 298 ophiolites, North Africa 29, 31, 33 orbital forcing 311 orbitolinids Gadvan Formation 353, 381 Galve sub-basin 114, 117, 119, 120, 121, 122, 123, 137 Kazhdumi Formation 355– 356 Ramales platform 324 Sarvak Formation 194, 203 Oxfordian Iberian Basin, epeiric carbonate ramp 93, 94, 95– 100, 106 see also intra-Oxfordian unconformity Oxfordian, Early, transgression 24, 25 oxygen isotopes Alps and Apennines 409 –410 Ramales platform 334–335 oyster buildups, Miocene, Ermenek Platform 275, 276, 278, 280, 282 Pabdeh Formation 223, 231 stratigraphy 234, 237, 238, 239 Padena outcrop 190, 192, 194, 200–206, 201 palaeoclimate, meteoric diagenesis 312 –313 Palaeogene, transgression 32, 31–34 palaeosol Mahil Formation 298–299, 299, 300 meteoritic diagenesis 301, 304, 305 Paleocene–Eocene Thermal Maximum 31 Palmyrides 44 tectonic setting 43, 45–46 Triassic carbonate-evaporite deposits 43, 47–61 hydrocarbon reservoirs 43, 44, 60– 61 Paquier Event 27 paragenesis Alps and Apennines 402, 404 Anaran Anticline 379 Pen˜agolosa sub-basin 115, 118 Peral outcrop, sequence stratigraphy 129– 130, 134 permeability 291 effect of meteoric diagenesis 310– 311

Permo-Triassic boundary 9 TSP 1 16– 18, 17 TST 1 16, 17 Persian Gulf, oolitic systems, comparision with Amellago ramp system 81, 84 platform-basin transition Galve sub-basin 113–139 Sarvak Formation 189– 214 pore systems meteoric diagenesis 291– 313 occlusion 294, 295– 296, 302, 305, 306 porosity effect of meteoric diagenesis 291, 294– 297, 303, 304, 305 petrophysical classes 310–311 mouldic 297, 298, 303, 304, 379, 381 Ramales platform 329, 330, 331, 333, 337– 340 vuggy 294–297, 300, 303, 305, 310 Anaran Anticline 366, 367, 371, 372, 384, 386 Portoles outcrop 129, 130, 134 Provence platform, sea level change and deposition 133, 135 –137 pyrite, Ramales platform 331 Ramales platform 317–340, 319 cementation 324–331, 338–339 depositional facies 321– 323, 324, 326, 327 diagenesis 324–331, 335–340 dissolution 331, 333, 338 –339 hydrothermal stage 339 porosity development 331, 333, 338–339 sequence stratigraphy 323–324 stable isotopes 334–335 stratigraphy 319, 320, 321 thermal modelling 332, 333– 334 uplift 339 Ranero limestones 320, 321– 324, 337– 338 reservoir modelling, Anaran Anticline 382, 384–387 Rhaetian, Palmyrides megacycle D 60 Rhaetian, Upper, Tropfbruch coral limestone 293, 294, 301 meteoritic diagenesis 302, 303, 304, 305 stable isotope data 307, 309 unconformities 301 rifting 18– 25 Rio Calera unit 320, 321 Rogenstein 293, 294 meteoric diagenesis 294– 297, 298, 311 stable isotope data 305, 308– 309 rudists Aptian, Galve sub-basin 120, 121, 130, 138 Cretaceous Oman epeiric carbonate platform 159 Sarvak Formation 194, 200, 203, 205 rugose topography, carbonate deposition model 14 Rupelian, Asmari Formation 236, 249, 253, 254, 255, 256 Salzedella sub-basin 115, 118 Samail Ophiolite 164 sandwaves Jabal Qusaibah 177, 178, 179 Jabal Shams 179, 180, 181

INDEX Sarvak Formation 187– 214, 191, 348 correlation with Natih Formation 207, 208, 209, 212 depositional environment 194– 197, 195 depositional sequence model 209– 211 dolomitization 348, 365, 366, 367, 368, 376 facies analysis 194– 197, 195 faults and joints 350, 351–353 lithostratigraphy 192, 193, 194 Lower 350, 357 dolomitization 376 stratigraphy 358–359 sequence heirarchy 206–207 sequence stratigraphy model 197– 206, 208 Bibi Seydan outcrop 197–200 Padena outcrop 191, 200–206, 201 stratigraphy 357–362 Upper dolomitization 348 stratigraphy 359–362 Scythian, Palmyrides megacycle A 49, 53, 55 sea level change channel incisions 165, 171 Dezful Embayment, Oligo-Miocene 248, 250, 252, 254 Ermenek Platform 283, 284, 285 Galve sub-basin 131, 133, 135, 135– 137 meteoric diagenesis 311– 312 Oman epeiric platform 157–160 Sarvak Formation 199– 200, 203– 204, 207, 209–210 see also eustacy sediment infill 309 Selli Event 25, 131, 159 Semail Ophiolite 33 Sergelu Formation 48, 51, 60 Serravalian, Ermenek Platform 271, 272, 286 shoal, tidal, Lower Dariyan Formation 373, 374 Shu’aiba Formation 26, 146, 147, 148 –150 biostratigraphy 137 correlation 150, 152 early Aptian OAE 159 sea level change and deposition 157– 160 seismic data 150, 151, 152 sequence stratigraphy 149 –150, 149 Siberian Traps LIP 21 siliciclastics, Asmari Formation 227, 229, 231 Simply Folded Belt, Zagros Mountains 344, 345 Sirte Basin 31, 33 Sopen˜a limestones 320, 321–323, 324, 326, 331, 337– 338 Soria Seaway 95, 100, 106 Spain Basque-Cantabrian Basin 317 –340 Galve sub-basin, Aptian platform-basin transition 113–139 Iberian Basin, Upper Jurassic epeiric carbonate ramp 89–108 stromatolites, Asmari Formation 235– 236 strontium isotope stratigraphy, Asmari Formation 240, 241, 242, 243– 245, 253 stylolitization Anaran Anticline 363–365, 380, 381 Mahil Formation 301, 303 Ramales platform 333 Tropfbruch coral limestone 306, 310

421

Surgah Formation 351, 360, 362 dolomitization 365 Syria, Triassic deposition 47 Syrian Arc tectonic event 29 Tang-e-Gurguda outcrop 225, 230 clinoforms 246, 250, 254 lithostratigraphy 236 Sr isotope stratigraphy 241, 242, 243 tectonostratigraphic phases (TSPs) 9, 11, 12– 13, 15–36 Tethys carbonate deposition 25–28 paleogeography, Triassic 46 tectonics 29–34 Tithonian, Iberian Basin, epeiric carbonate ramp 92, 93, 94, 96, 97, 101, 103, 104, 105, 107– 108 Toarcian, Late, unconformity, TSPs 11, 12, 20, 21, 22 Toarcian transgression 21, 28 total organic matter (TOC), Oman epeiric carbonate platform 158 transgression, basins 11, 14–15 transgressive systems tracts (TSTs) 15– 36 Galve sub-basin 127, 129 –131 transgressive-regressive cycles, Jurassic ramps 95, 100 Triassic Arabian Plate, deposition 46–47 meteoric diagenesis, karst unconformities 291–313 Palmyrides, carbonate-evaporite deposits 43, 47– 61 Triassic, Late, transgression 18, 19 Triassic, Lower Calvo¨rde Formation 293, 297 stable isotope data 307, 308– 309 Triassic, Middle Mahil Formation 293, 298–305 stable isotope data 305, 309 transgression 17, 21 unconformity, TSPs 11, 12, 17, 18–20 Triassic– Jurassic boundary, extinction 22 Tropfbruch coral limestone 293, 294, 301, 302 meteoric diagenesis 302, 303, 304, 305, 311 stable isotope data 307, 309 Turkey, Ermenek Platform 265–288 Turonian Early, transgression 27, 28 Middle, unconformity, TSPs 11, 13, 28, 30, 27– 31 see also Cenomanian– Turonian unconformities karst Dariyan Formation 354– 355 meteoric diagenesis 291 Mahil Formation 299, 305 Ramales platform 323, 337 Tropfbruch coral limestone 301, 305–306 TSPs 9, 11 see also names of individual stages Urgonian carbonate platforms 114, 118, 135, 138, 317, 318 USA, fracture-controlled dolomite 391 –392 Utrillas Formation 116, 118, 123

422

INDEX

Valmaseda Formation 320, 321, 338 Vercors platform interior domain, sea level change and deposition 133, 135– 137 Villarroya de los Pinares Formation 116, 117, 118 biostratigraphy 122, 123, 124 carbon isotope excursion 125 depositional model 120 sequence stratigraphy 129, 130, 131, 134

Well-13 Asmari Formation biostratigraphy 237 lithostratigraphy 236, 237, 240 West Caicos, oolitic system, comparison with Amellago ramp system 84

Wadi Aday, Mahil Formation 298–305 Wadi Mi’Aidin, tidal channel fill 172 Well-2, Asmari, Pabdeh, Jahrum formations, lithostratigraphy 239, 240 Well-5, Asmari Formation, lithostratigraphy 238, 240 Well-9, Asmari Formation, Sr isotope stratigraphy 242, 253 Well-12, Asmari Formation, Sr isotope stratigraphy 242

Zagros fold belt, fracture-controlled dolomite 388–389 Zagros Foredeep 35 Zagros Mountains 344, 345 fracture-controlled dolomitization 343–393 geology 189– 190, 220, 221, 222 zebra fabrics, Anaran Anticline 371 –374, 373, 374, 377 zinc ore, Ramales platform 331

Yenimahalle Formation 270

This volume contains a collection of stratigraphic and diagenetic case studies of Mesozoic and Cenozoic carbonate sequences from the Tethyan realm. High levels of industry and academic interest in the region have generated numerous multi-disciplinary studies of these sequences, a selection of which are presented in this volume. The studies presented are based on both comprehensive subsurface datasets from important hydrocarbon-bearing strata of the Middle East and the excellent surface exposures in the region of interest. The studies presented in this volume may serve as suitable starting points in the development of age and architecture specific carbonate reference models. Such models can form the basis of internally consistent models for carbonate deposition, sequence development and reservoir performance. Ideally such models, suitably scaled, will be equally applicable to academic studies, the exploration and development phases of the field life cycle and in the prediction of future reservoir performance.